N₂O, a greenhouse gas, has an atmospheric lifetime of about 120 years, so any changes recorded in ice cores reflect global changes, thought to mainly reflect increased application of agricultural fertilizers (Wolff 2013). N₂O concentrations recorded in air bubbles within the ice have risen by about 20% in the last 200 years (see Law Dome example below) and are at levels higher than recorded for the last 800 ka (Wolff 2013). In contrast, ammonium (NH₄⁺), mainly deposited as aerosol, and nitrate (NO₃⁻), deposited as an aerosol or as nitric acid, have much shorter atmospheric residence times, and changes in their concentrations reflect local changes in source areas (Wolff 2013). Mid-latitude glaciers, such as Mount Muztagata in China, show an approximately three-fold increase in ammonium concentrations between 1960 and 1990 CE, probably as a consequence of agricultural emissions (Zhao et al. 2011).

However, there is no significant recent trend at either pole (Wolff 2013). Whereas ammonium is permanently fixed in the snow, nitrate can be re-emitted back to the atmosphere (Fischer et al. 1998, Wolff 2013). Increased nitrate concentrations evident in the Northern Hemisphere (Figure 25) are mainly due to NOₓ emissions from fossil fuel combustion (Wolff 2013, 2014), including automobiles and coal-fired power stations. In Greenland ice the main phase of nitrate increase, which has risen by a factor of 2–3, occurred from 1950 to 1980 CE (Fischer et al. 1998, Wolff 2013); by 1980 CE levels were higher than over the past 100 kyr (Wolff 2014). In the French Alps, at the Col du Dôme, high snow accumulation rates allow distinction between summer and winter snow, with higher summer nitrate values showing a fivefold increase during the 20th century, most markedly between 1960 and 1980 CE and sourced by emissions within 1000 km of the glacier (Preunkert et al. 2003). However, nitrate does not provide a consistent global signal. In central Asia, at Mount Muztagata, the steep rise in nitrate concentrations, which increased
by a factor of two, started after the late-1970s and had already peaked by 2000 CE (Zhao et
al. 2011), whereas no significant increase occurs then in Antarctic ice (Wolff 2013).

A shallow (100 m) GISP2 ice core from Summit (Greenland), ranging in age from 1718 to 2006 CE, shows a clear decreasing δ¹⁵N trend in atmospheric nitrate, from pre-industrial values around 11 to -1‰ in the early 21st century (Hastings et al. 2009). The δ¹⁵N trend began about 1850 CE, whereas the increase in nitrate from the same core rises in about 1890 CE (Hastings et al. 2009). The greatest rate of change in δ¹⁵N occurred from 1950 to 1980 CE, coincident with a rapid rise in fossil fuel emissions (Hastings et al. 2009).

During pre-industrial times, about half the lead in the troposphere was associated with soil
dust and half with volcanic gases (Patterson and Settle 1987). Pb, As, Sb, Bi, Cu, and Zn
enrichments in Arctic ice cores started at around 3300-3000 yr BP, with an additional
significant Roman phase (Hong et al. 1994, Krachler et al. 2008, 2009). Further increases of
Pb, Tl, and Cd in response to the Industrial Revolution are more significant, with elevated
concentrations from the 1860s, accelerated concentrations from 1887 to 1889 CE and peaks
in the early 20th century (TI in 1911-1915 CE, Cd in 1906 CE, Pb in 1915 CE) (McConnell and
Edwards 2008, Figure 26). This early peak is considered to reflect increased coal burning in
North America and Europe (McConnell and Edwards 2008). However, the greatest rises in Pb
concentrations are from the 1950s, peaking in the 1960s, associated with emissions of Pb
used in alkyl-leaded gasoline (Murozumi et al. 1969, Boutron et al. 1991, McConnell and
Edwards 2008). TI shows no equivalent mid-20th century peak, whereas Cd has a 1946 CE
peak concentration (Figure 26). At its peak, Pb levels in Greenland ice were above Holocene
background levels by a factor of 200 and exceeded concentrations seen during the last
glacial maximum (Murozumi et al. 1969); in Antarctic ice they were a factor of five above
pre-Industrial levels (Patterson and Settle 1987, Wolff and Suttie 1994). Use of lead additives in gasoline became limited in the USA and other Northern Hemisphere countries from about 1970 CE, and concentrations decreased in Greenland by a factor of 7.5 between 1970 and 1990 CE, with Cd (Figure 26) and Zn sourced from industrial processes showing a decrease by a factor of 2.5 over the same interval (Boutron et al. 1991, McConnell and Edwards 2008). In Antarctica, by 1920 CE anthropogenic Pb (mainly from metal processing) dominated concentrations that were 2 to 4 times higher than pre-Industrial levels (Wolff and Suttie 1994). Lower concentrations in the 1930s to mid-1940s are followed by a marked upturn in concentrations by a factor of >2 from 1950 to 1980 CE, with subsequent decline (Wolff and Suttie 1994). The first use of unleaded gasoline in many Southern Hemisphere countries started as late as the 1980s and 1990s (Wolff and Suttie 1994). Because most combustion of gasoline occurred in the Northern Hemisphere, it is in Greenland that the rise and subsequent fall of Pb signals is greatest. Lead aerosols have a 10-day residence time in the atmosphere resulting in little inter-hemispheric interchange (Patterson and Settle 1987).

<figure><img src="image_37.png" /></figure>
**Figure 26.** Concentrations (5-year running means) in Greenland ACT2 ice core from 1772 to 2003 CE of sulphur (NssS), thallium (TI), cadmium (Cd) and lead (Pb), compared to black carbon (BC), from McConnell and Edwards (2008). Sulphur peaks in 1817 and 1883 CE relate to the Tambora and Krakatoa volcanic eruptions, respectively. © (2008) National Academy of Sciences, U.S.A.

Atmospheric Hg contamination in a glacier in Wyoming (USA) from ~1720 to 1993 CE shows a twenty-fold increase from pre-1840 CE levels to a ~1984 CE peak (Schuster et al. 2002). Early spikes include both volcanic sources from Tambora (1815 CE) and Krakatoa (1883 CE), showing cross-hemispheric transfer of Hg, and an anthropogenic source with a double peak and elevated values from ~1850 to 1884 CE associated with the California Gold Rush in western USA (Schuster et al. 2002).

In Greenland, black carbon is attributed by McConnell and Edwards (2008) as: Preindustrial (1772–1860 CE), coal-dominated industrial (1860–1940 CE), and oil-dominated industrial (1940–2003 CE), with peak concentrations in the first decade of the 20th century (Figure 26). However, peak combustion of coal is later in Asia. Black carbon and organic carbon concentrations in Tibetan ice cores show a marked high during the 1950s and 1960s, attributed to the inflows of industrial emissions, mainly from Europe, with a decline during the 1970s and 1980s (Xu et al. 2009). Glaciers in the southern part of the plateau show black and organic carbon increases after the 1990s, from a southern source (Xu et al. 2009). Increased black carbon on glacial ice reduces surface reflectivity and increases melting (Xu et al. 2009), potentially causing significant hiatuses in modern laminae.

Microplastics have been reported in Arctic sea-ice (Obbard et al. 2014) from contaminated surface sea-water that has been frozen. However, there is no record of microplastics in polar
continental ice sheets or glaciers. The potential for airborne transfer of microfibres (Dris et al. 2016) suggests that microplastic fallout in glacial ice might occur, especially where glaciers are close to urban areas.

Table 10 summarises the key advantages and disadvantages of having a potential GSSP within glacial ice, as is the case for the base of the Holocene. The Law Dome ice core from Antarctica has a high-resolution multi-proxy record for the last 1000–2000 years and shows many characteristics that make it suitable for hosting a potential GSSP. Other similarly suitable ice cores are available elsewhere in both Antarctica and Greenland (Figure 23), although many of the sites are unsuitable because the rate of accumulation is too low (e.g. Dome C, Antarctica).

<table><caption>**Table 10.** Reasons for and against using an ice core as a potential host for a GSSP.</caption><thead><tr><th>For:</th><th>Against:</th></tr></thead><tbody><tr><td><ul><li>Annual layers, high resolution</li><li>Extensive distribution, but only Greenland and Antarctic ice sheets likely to persist</li><li>Independent dating using <sup>210</sup>Pb, but not <sup>14</sup>C</li><li>Major volcanic eruption events provide clear S spikes</li><li>SO<sub>4</sub><sup>2-</sup> clear spike starts mid-20<sup>th</sup> century, but diachronous</li><li>NO<sub>3</sub><sup>-</sup> increase and &delta;<sup>15</sup>N depletion from ~1950 CE</li><li>Increased lead accumulation from 1950s, especially from gasoline source</li><li>Increased Zn, Cd, and Cu concentrations</li><li>Radionuclides, clearly resolved bomb spike</li><li>Holocene GSSP precedent in Greenland</li></ul></td><td><ul><li>Surficial laminae (recent decades) are more prone to melting by global warming than deeper laminae (Holocene), especially in alpine glaciers and coastal ice sheets</li><li>Alpine glaciers melting at rapid rate</li><li>Lag between age of ice and younger age of air bubbles (100 years in Greenland, possibly up to 1000 years in Antarctica)</li><li>&delta;D and &delta;<sup>18</sup>O no significant variation across mid-20<sup>th</sup> century</li><li>Increased dust flux not consistently evident</li><li>Mid-20<sup>th</sup> century increased CO<sub>2</sub> and CH<sub>4</sub> concentrations and &delta;<sup>13</sup>C depletion signal not fixed in air bubbles</li><li>Core difficult/expensive to procure and to store safely</li></ul></td></tr></tbody></table>
3.7.1 Law Dome Ice Core, East Antarctica

A high-precision 1000–2000 year history of CO2, CH4, N2O and δ13C has been acquired from the Law Dome site in East Antarctica (Etheridge et al. 1996, Francey et al. 1999, Ferretti et al. 2005, MacFarling Meure et al. 2006, Rubino et al. 2013). As regards hosting a potential GSSP site, the high snow accumulation rates, low temperatures, and small quantities of impurities at Law Dome make this record particularly suitable, especially for recognising a high-resolution δ13C signal (Rubino et al. 2013). The Law Dome site comprises three cores drilled between 1987 and 1993 CE, (Etheridge et al. 1996, Rubino et al. 2013), dated by counting annual layers in oxygen isotope ratios (δ18O), ice electroconductivity measurements and hydrogen peroxide (H2O2) concentrations (Etheridge et al. 1996), with a gas-age/ice-age difference of between ~31 and ~55 years for CO2 (Rubino et al. 2013).

Between 1000 and 1600 CE, the CO₂ concentration varies between 278 and 284 ppm (Figure 27a), and between 1000 and 1500 CE, the δ¹³C values remained constant at about -6.55‰ (Rubino et al. 2013). The brief decrease in concentration of CO₂ in the Law Dome ice core around 1600 CE (Etheridge et al. 1996, MacFarling Meure et al. 2006) was used by Lewis and Maslin (2015) as their “Orbis Event” which they regarded as sufficiently important to be a candidate for defining the beginning of the Anthropocene. Analysis of changes in atmospheric carbonyl sulphide concentration, which is linked to changes in gross primary production of terrestrial ecosystems, shows that temperature change, rather than vegetation re-growth, was the main cause of the increased terrestrial storage and hence drop in atmospheric CO₂ around 1600 CE (Rubino et al. 2016). After a recovery, CO₂ maintained comparatively low concentrations and elevated δ¹³C through to 1750 CE (Rubino et al. 2013). Over the past 200 years, atmospheric CO₂ concentrations at Law Dome have increased by 29% (MacFarling Meure et al. 2006). Following the Industrial Revolution,
Rubino et al. (2013) show two inflexion points in the 20th century $\delta^{13}$C signal: from 1915 CE, and from 1955 CE. There is stabilization of atmospheric CO$_2$ concentrations at 310–312 ppm from ~1940 to 1955 CE (Figure 25) (MacFarling Meure et al. 2006). The subsequent increase appears to coincide with increase in fossil fuel use at about 1950 CE (Rubino et al. 2013). The Francey et al. (1999) high-precision record of atmospheric $\delta^{13}$C based on Antarctic ice cores shows two straight segments: one between 1850 and 1961 CE, and a steeper segment between 1962 and 1980 CE.

<figure><img src="image_38.png" /></figure>

<figure><img src="image_39.png" /></figure>
<figure><img src="image_40.png" /><figcaption>**Figure 27.** a) Law Dome ice core and firn records for CO₂ concentration and δ¹³C for atmospheric CO₂ for the past 1000 years (from Rubino et al. 2013); b) Law Dome ice core and firn records for CH₄ concentration and δ¹³C for atmospheric CH₄ for the past 1000 years (from Ferretti et al. 2005); c) N₂O concentrations in the Law Dome ice core over the past 2000 years and 200 years (Wolff 2013).</figcaption></figure>

The Law Dome ice core records also show atmospheric methane (CH₄) concentrations ranging from 590 to 760 parts per billion (ppb) through much of the Holocene, up to 1700 CE, with as little as ~55 ppb difference from 1000 to 1700 CE (Ferretti et al. 2005). This is followed by an unprecedented increase to 1700 ppb by 2004 CE (Ferretti et al. 2005, Figure 27a), an increase of 150% over 200 years (MacFarling Meure et al. 2006). Growth rates in CH₄ decreased during ~1940–1955 CE (MacFarling Meure et al. 2006). The δ¹³C curve for CH₄ shows a marked decrease of ~2‰ from ~1000 to 1700 CE, perhaps in response to reduced biomass burning, and a subsequent abrupt rise from ~1875 CE to the present value of ~-47‰, reflecting increasing pyrogenic emissions (Ferretti et al. 2005, Figure 27a).
Over the past 200 years, atmospheric N₂O concentrations at Law Dome increased by 21% (MacFarling Meure et al. 2006). N₂O concentrations ranged from 260 to 270 ppb between 1 and 1850 CE, followed by a steady rise to 290 ppb by 1950 CE, from when there was a rapid rise to 320 ppb by 2000 CE (Wolff 2013, Figure 27b). The δ¹⁵N signal in firn ice at Law Dome is more scattered than would be expected, most probably due to random analytical noise (Rubino et al. 2013).

## 3.8 Speleothems

Annually laminated calcareous speleothems (typically stalagmites) occur within natural cave systems, typical of karst environments, or within artificial tunnels where either the adjacent bedrock or the degradation of mortar in concrete linings of the tunnel contributes the calcium carbonate. Karst landscapes occur across extensive parts of Eurasia, Australasia and the Americas, in total about 10% of the Earth's land mass (Figure 28). There is precedent in using speleothems for the location of a GSSP, as Walker et al. (2012) used a stalagmite from Mawmluh Cave, India, as the proposed GSSP for the Upper Holocene Subseries (= proposed Meghalayan Stage) at 4.2 ka.

In addition to lamina counting, the age of older speleothems is typically determined by ²³⁴U-²³⁰Th dating, although this is unsuitable for carbonates precipitated over the past century. Independent dating using ²¹⁰Pb-²²⁶Ra has been undertaken at a few locations (e.g. by Condomines and Rihs 2006 at St. Nectaire, France, and Bonotto et al. 2012 at Santana Cave in Brazil). Radiocarbon has quite commonly been used to confirm an age model from annual laminae, but does not provide annually resolved dates. If Pu is chosen as a primary tool for
correlation, the lack of Pu isotope determination would rule out a speleothem as GSSP for
the Anthropocene (Fairchild 2017), but one could be chosen as an auxiliary site.

<figure><img src="image_41.png" /><figcaption>**Figure 28.** Extent of carbonate outcrops present across global land mass by Ulrichstill https://commons.wikimedia.org/w/index.php?curid=9412430 Created 6th February 2010. Karst landscapes and speleothems occur within these regions, and locations mentioned in the text are indicated.</figcaption></figure>

Natural speleothem growth is typically slow, on the order of tens to hundreds of micrometres per year. In contrast, speleothems related to hyperalkaline groundwaters, commonly from anthropogenic sources, can grow at a rate of 10 mm per year (Baker et al. 1998), hence potentially permitting fine-scale resolution of environmental changes, although the chemistry will be strongly kinetically modified (Hartland et al. 2014, Newton et al. 2015).

²¹⁰Pb dating of stalactites in Santana Cave (Brazil) shows a longitudinal growth rate of 1.3 mm yr⁻¹ and lateral rate of 0.01 mm yr⁻¹ (Bonotto et al. 2012). A study of hydrothermal
stalagmites at St Nectaire (France) using $^{210}$Pb/$^{226}$Ra ratios showed growth rates varying from 5.3 ± 0.5 mm to 2.6 ± 0.2 mm yr$^{-1}$ (Condomines and Rihs 2006).

Speleothems record changing environmental conditions in the atmosphere, soil and ecosystem through geochemical signals transmitted from the ground surface, through rock, to the subterranean void. The processes of speleothem formation and their chemistry are detailed by Fairchild and Baker (2012), and their relevance to recognition of Anthropocene signals is given in Fairchild and Frisia (2014) and Fairchild (2017). Principal signals that may be used to recognise a potential Holocene–Anthropocene boundary (Table 2) include shifts in atmospheric $^{14}$C and sulphur (sulphate concentration and $\delta^{34}$S), which relate to global-scale changes to atmospheric chemistry (Fairchild and Frisia 2014, Figure 29). Local signals include: 1) variations in growth rates of laminae and $\delta^{18}$O which relate to air temperature and humidity, 2) $\delta^{13}$C as an indicator of deforestation and/or introduction of C4 plants, 3) development of biomarkers such as a reduction in the ratio of C$_{27}$/C$_{31}$ n-alkanes and increase in n-alkanols as observed at Mechara (Ethiopia) reflecting the local introduction of agriculture about 1935 CE (Blyth et al. 2016), and 4) shifts in trace elements and isotope ratios (Fairchild 2017). Although of local source and probably diachronous in nature, such signals may link to global patterns of environmental change and so contribute to characterising a potential GSSP site. For example, Pb anomalies have been recorded in speleothems in Somerset, UK near mines active in pre-Roman, Roman and mediaeval times (McFarlane et al. 2013). Sulphur pollution from coal combustion increased during the Industrial Revolution and culminated in peak S concentrations and depleted $\delta^{34}$S values in the mid- to late-20th century in developed countries (Wynn et al. 2010, Figure 29d). However, speleothems are insensitive to the significant increases in atmospheric CO$_{2}$ concentrations and depletion in $\delta^{13}$C over the past century (Fairchild and Frisia 2014).
Changes in the nitrogen cycle via N abundance or isotopic fractionation related to fertiliser use have not been investigated in speleothems (Fairchild 2017). Phosphorus is strongly modulated by seasonal vegetational die-back and changed infiltration rates (Fairchild and Frisia 2014) rather than by fertiliser application.

Transport of solutes via plants, soil and rock commonly results in signals being variably attenuated and delayed (Fairchild and Frisia 2014, Fairchild 2017). For example, speleothems in SW France and Gibraltar show ¹⁴C concentrations with differing degrees of attenuation and delay of the peak signal, but the initial rise was consistent at 1958 CE and coincides with changes in atmospheric concentrations (Fairchild and Frisia 2014). The timing of onset of the ¹⁴C rise in speleothems is typically within 1–2 years of the atmospheric change (Genty and Massault 1999).

Table 11 summarises the key advantages and disadvantages of such a potential GSSP for the Anthropocene. As examples, the Ernesto Cave in Italy has multi-proxy studies of speleothem laminae within a natural cave system, but urban speleothems may show greater tuning to modified anthropogenic signals.

<table><thead><tr><th>For:</th><th>Against:</th></tr></thead><tbody><tr><td><ul><li>Annual lamination</li><li>Undisturbed locations</li><li>Independent ¹⁴C markers</li><li>S, δ³⁴S, ¹⁴C reflect atmospheric composition, but modified by soil-ecosystem</li><li>Solid samples easier to handle than unconsolidated sediment</li><li>Precedent of Mawmluh Cave, India, for the proposed Upper Holocene GSSP</li></ul></td><td><ul><li>The atmospheric Suess effect on δ¹³C is smaller than fractionation effects in the soil-ecosystem so an inflection corresponding to atmospheric change is not normally present</li><li>Although S, δ³⁴S, ¹⁴C may show initial rise as atmosphere changes, the peak is delayed because of storage in soil/ecosystem</li><li>Changes in trace pollutant metals (e.g. Pb, Zn, REE) are strongly linked to organic matter from soil and increases tend to reflect soil disturbance/deforestation effects</li><li>Not able to reliably detect signals usable elsewhere, e.g. Pu, anthropogenic detritus</li></ul></td></tr></tbody></table>
**Table 11.** Reasons for and against using a speleothem as a potential host for a GSSP.

3.8.1 Ernesto Cave, Italy

Ernesto Cave, northern Italy, closely studied for >20 years, has supplied various proxy data series from annually-laminated speleothems independently dated using U-Th series (Frisia et al. 2003). These show a strong annual signal of fluorescent laminae associated with a range of colloid-transported elements (Fairchild and Frisia 2014). A pronounced change is recognised since about 1840 CE, from open and impurity-rich crystals within relatively thin laminae to thicker laminae of clean calcite (Figure 29a), representing higher growth rates, closely correlated with temperature increases (Frisia et al. 2003, Fairchild and Frisia 2014) associated with the local ending of the Little Ice Age. This pattern broadly coincides with the initial increase in sulphate concentrations within speleothems in this cave at about 1880 CE (Frisia et al. 2005), associated with the initiation of atmospheric pollution in response to the Industrial Revolution.

No appreciable change in stable isotopes has occurred across the mid-20th century (Figure 29b), the most significant change in δ¹³C being a marked positive excursion at ~1840 CE (Scholz et al. 2012).
<figure><img src="image_42.png" /></figure>

<figure><img src="image_43.png" /></figure>

<figure><img src="image_44.png" /></figure>
<figure><img src="image_45.png" /><figcaption>**Figure 29.** Ernesto Cave, Italy: a) Age model based on lamina counting related to local mean air temperature (Frisia et al. 2003); b) δ¹⁸O and δ¹³C profiles (Scholz et al. 2012); c) Radiocarbon profile and comparable European atmospheric emissions (Fohlmeister et al. 2011); d) S concentration and δ³⁴S (Frisia et al. 2005, Wynn et al. 2010).</figcaption></figure>

The ¹⁴C bomb-spike has been recorded from the cave (Figure 29c), but with a decade lag compared with atmospheric signals (Fohlmeister et al. 2011). Records of sulphur loading and lowering of δ³⁴S in both speleothem and tree rings show a similar lag in the speleothems of about 15–20 years (Figure 29c), such that in 2000 CE the speleothem was still recording peak atmospheric loading (Wynn et al. 2010, 2014) that was locally experienced in the 1980s
(Fohlmeister et al. 2011), but which is now declining (Borsato et al. 2015). Trace element concentrations increased greatly during the early-20th century, a time of deforestation in the area (Borsato et al. 2007).

### 3.8.2 Urban speleothem, Paris

Rises in several trace metals and rare earth elements are found in urban speleothems in an underground aqueduct beneath Paris that have grown over 300 years with a record of two laminae per year, independently dated using U-Th series (Pons-Branchu et al. 2015). Two distinct phases of heavy metal pollution are recognised: a) 18th century metal contamination resulting from use of urban waste (night soil) as fertilizer; and b) since 1900 CE elevated Pb, Mn, V, Cu, Cd and Al concentrations, attributed to urbanisation, with a more marked increase since 1960–1970 CE (Pons-Branchu et al. 2015). $^{206}$Pb/$^{207}$Pb isotope ratios show no consistent pattern in these speleothems. The dominant value of $1.181 \pm 0.003$, is typical of values associated with lead from northern Europe (Pons-Branchu et al. 2015), though one speleothem showed a change after the mid-19th century, suggesting a new contribution from coal burning or from Pb from Spain. Pb ratios typical of leaded gasoline are not discriminated (Pons-Branchu et al. 2015).

### 3.9 Trees

Forest coverage in 2005 CE was about 39 million km$^2$, or ~30% of the world's land area (FAO 2006) and there are an estimated 3 trillion trees (Crowther et al. 2015). Forests are extensive on all continents except Antarctica, and include the boreal forests of North America and Eurasia and tropical forests of South America, central Africa and south-east Asia (Figure 30). Tropical trees from aseasonal tropical regions lack reliable annual rings, so trees from
temperate or boreal environments would be most suitable for a GSSP as the tree rings
themselves can provide a precise annually-resolved chronology.

<figure><img src="image_46.png" /><figcaption>**Figure 30.** Forest map of the world with key locations mentioned in the text. © (2006)FAO.]</figcaption></figure>

Tree rings provide a perfect annually-resolved chronology, and within this an annually resolved archive of environmental change, most notably palaeoclimate, at local to global scales. This signal can be evident in the physical properties of the tree ring (width, density, reflectance) or chemical indicators such as the stable isotopic ratios of carbon (sourced from atmospheric CO₂), and hydrogen and oxygen (sourced from precipitation, soil or ground water) (Fritts 1976, Speer 2010, McCarroll and Loader 2004). In many species, the tree rings may be further divided into early wood (formed during the spring) and late wood (formed during the summer, primarily from photosynthates fixed during the growth year). Rings from living trees can provide environmental information spanning both the mid-20th century putative boundary for the Anthropocene and the late-18th century transition from pre-
industrial times through the development of the Industrial Revolution. Similar records can also be found in preserved dead trees that accumulated in bogs or lakes, and also in building timbers. Where trees co-existed in time they will co-record a shared environmental signal that may therefore be cross-dated to assign chronology or to provide independent replication between trees or across study regions.

Table 12 summarises the key advantages and disadvantages of having a potential GSSP within tree rings. Proxy signals associated with palaeoclimate, stable carbon isotopes, sulphur concentrations and isotopic ratios, radiocarbon fallout and heavy metal concentrations are discussed below.

<table><caption>**Table 12.** Reasons for and against using a tree as a potential host for a GSSP.</caption><thead><tr><th>For:</th><th>Against:</th></tr></thead><tbody><tr><td><ul><li>Global expression and distribution; widespread terrestrial extent</li><li>Bi-annual layers</li><li>Precise dating through dendrochronology</li><li>Independent <sup>14</sup>C dating possible</li><li>Environmental signals preserved in deadwood</li><li>δ<sup>18</sup>O and δD: record source water, which contains a climate signal</li><li>δ<sup>13</sup>C is highly sensitive to the Suess effect on atmospheric CO<sub>2</sub></li><li>Sulphur and δ<sup>34</sup>S signal commencing mid-20<sup>th</sup> century</li><li><sup>14</sup>C bomb spike clearly defined</li><li>High potential for independent replication</li><li>Tangible, stable and archivable</li></ul></td><td><ul><li>Growth ring width may not be suitable for picking mid-20<sup>th</sup> century climate signal (due to local effects or 'divergence')</li><li>Complex record of local signals in physical and chemical proxies</li><li>δ<sup>13</sup>C can be affected by local modifications</li><li>Heavy metal concentrations may be unsuitable proxy for pollution events due to radial mobility</li><li>No precedent of using a living/deceased organism for GSSP</li></ul></td></tr></tbody></table>
### 3.9.1 Tree rings and palaeoclimate signals

Dendrochronology relies upon the counting and statistical synchronisation of annual changes in tree ring widths, and their assignment to a calendrical or relative timescale. This is possible because the variability in tree growth driving this signal is related to regional climatic changes. Trees growing at the same time in history, and experiencing a similar climate history, will in the absence of external disturbances yield similar patterns of wide and narrow growth rings that may be cross-dated by dendrochronology to develop a precise timescale. Due to the physical increase in the diameter of a tree as it grows, tree-rings show a characteristic decline in width with increasing tree age (size). Where a climatic signal is to be extracted from the ring width data, statistical techniques are required to remove this growth or age-related trend. This statistical processing may cause climate signals of frequency longer or equal to the age-span of the tree to be lost, reducing the value of tree-ring widths in recording long-term climate change (Cook et al. 1995, Esper et al. 2002). However, this effect can be significantly reduced/removed by using more complex detrending protocols and numerous multi-centennial records of temperature changes from 800 to 1990 CE have been determined from high-elevation and high-latitude locations across the Northern Hemisphere using Regional Curve-based Standardization techniques (Esper et al. 2002, Melvin and Briffa 2008). These data contain evidence for a Northern Hemisphere Medieval Climatic Anomaly (MCA) between about 900 and 1200 CE, and for below average temperatures typical of the Little Ice Age (LIA) from 1200 to 1850 CE, and for warming since 1850 CE comparable in scale to that of the MCA (Wilson et al. 2016, Esper et al. 2002, Figure 31a).

The suitability of using tree rings to establish the temperature profile post-1950 CE has been questioned due to a conspicuous increase in the statistical uncertainty in reconstructions
after 1975 CE, which may reflect an anomalous reduction in growth performance at some
high northern latitude conifer sites, but is also often coincident with a decrease in the
replication of the composite record (Esper et al. 2002). This "divergence" between growth-
based reconstructions and observed climate has been the subject of significant attention by
dendroclimatologists and many potential causes, from unbalanced replication and signal
processing effects to changes in seasonality and CO2 fertilisation, have been considered
(D'Arrigo et al. 2007). Whilst this may reduce current confidence in the use of ringwidth-
based reconstructions to define climatic changes attributed to the start of the
Anthropocene, maximum late wood density appears to be less impacted upon by such
divergence issues and density-based reconstructions of past climate remain some of the best
records of past climatic variability currently available to climate scientists. There are few
palaeoclimate proxies other than tree-rings capable of providing the same precisely dated,
annually-resolved palaeoclimate reconstructions with the same capacity for independent
verification through replication, capable of being objectively calibrated against instrumental
observations and independently verified.
<figure><img src="image_47.png" /></figure>

<figure><img src="image_48.png" /><figcaption>**Figure 31.** a) The 20-year smoothed Northern Hemisphere extratropics reconstruction of radial stem productivity in high elevation and high latitude forest environments since 800 CE (black) and two-tailed 95% bootstrap confidence intervals (blue) (from Esper et al. 2002); b) δ¹³C variability from Loader et al. (2013a) for the period 1500–2008 CE measured in tree-ring cellulose for a composite tree ring stable isotope chronology developed using *Pinus sylvestris* trees from northern Fennoscandia. Fine line represents annually-resolved δ¹³C variability, thick solid line presents the</figcaption></figure>
annual data smoothed with a centrally-weighted 51-year moving average. Dashed line represents the mean δ¹³C value for the “pre-industrial” period 1500–1799 CE. Mean annual replication for the record is >13 trees. Analytical precision σn-1 = 0.12 per mille n=951 which compares favourably with the analytical precision of the method typically reported (σn-1 = 0.10 n=10) (Boettger et al. 2007, McCarroll and Loader 2004, Loader et al. 2013a).

3.9.2 Stable carbon isotopic signal

Stable isotope variability in tree rings appears to be less affected by non-climatic age-related
tendencies (Young et al. 2011). Long-term trends of stable carbon isotopic data in tree rings
may therefore mark the start of the Anthropocene in the tree ring archive. The stable carbon
isotopes in tree rings record primarily the stable isotopic composition of the source CO2
sampled by the tree during photosynthesis. This signal is further imprinted upon by the
tree's physiological response to climatic change (the balance between stomatal conductance
and photosynthetic rate), photosynthetic and respirative processes (Farquhar et al. 1982,
McCarroll and Loader 2004, and references therein). Inter-annually, δ¹³C values will vary in
response to changes in irradiance, soil moisture, precipitation and relative humidity and this
has allowed the tree rings to be used as a record of past climate variability. As a biological
system, tree ring stable carbon isotopic ratios will also vary depending upon the species of
tree (angiosperm versus gymnosperm) and their individual environmental history. In
addition, factors such as leaf morphology, ecological amplitude, pollution, forest
management history, disturbance and nutrient availability, can all influence the resulting
tree ring isotopic record, but at all times, the basic physiological response of the tree will
remain the same. This natural inter-tree variability has been characterised in resampling
experiments (Loader et al. 2013b), and although extreme cases have been reported (Li et al.
2005), natural intra-site variability falls within reasonable limits defined by current analytical uncertainties.

When viewed over a long (centennial-millennial) timescale, trees have the capacity to record and contextualise the unprecedented anthropogenic changes in atmospheric $\delta^{13}C$ associated with global industrialisation and land-use change. By combining records from different geographical regions, location-specific year-to-year climatic variability and the effect of random disturbances may be reduced to yield a regionally relevant record of carbon isotopic variability in atmospheric CO$_2$. A reduction of atmospheric $\delta^{13}C$ values by ~2‰ has occurred since the start of the global Industrial Revolution ~1820 CE (the $\delta^{13}C$ Suess effect), which is among the dates proposed as a start for the Anthropocene (Figure 31b). This effect is observed in tree rings worldwide and is also clearly observable in ringless trees from the aseasonal tropics (Loader et al. 2011).

The $\delta^{13}C$ Suess effect is commonly removed in tree ring carbon isotope series through mathematical detrending without loss of low-frequency climatic information (McCarroll and Loader 2004). This makes the data suitable for numerical calibration against instrumental climate data, but prevents its use as an arbiter of atmospheric $\delta^{13}C$ changes. To assess the potential base of the Anthropocene using $\delta^{13}C$ changes, raw carbon isotope data are required.

A further consideration when assessing the candidature of tree ring $\delta^{13}C$ as a marker for the Anthropocene is the influence of increased atmospheric CO$_2$ concentration on photosynthesis. This "fertilisation" or "greening" effect has been widely reported and in the context of stable carbon isotopic variability is seen as a change in the intrinsic water-use
efficiency of the tree (Seibt et al. 2008, Saurer et al. 2014, Frank et al. 2015). Methods have been developed to quantify and to correct for this effect, which will vary from tree to tree (Treydte et al. 2009, McCarroll et al. 2009). Several studies have also identified a limit to the degree to which trees can adapt to higher atmospheric CO₂ concentrations. This “fertilisation” effect does not negate the use of δ¹³C as a potential marker, but it should be taken into consideration.

A large-scale/global composite of global carbon isotope time-series from tree rings could help locate the boundary of the Anthropocene, as individual records are affected by local climate, forest disturbance, or the local impact of atmospheric pollution on stomatal conductance and disease (McCarroll and Loader 2004 and references therein, Bukata and Kyser 2007, Treydte et al. 2007, 2009, Loader et al. 2013b, Boettger et al. 2014, Saurer et al. 2014, Frank et al. 2015), all of which could potentially obscure the true expression of the global δ¹³C Suess effect in tree rings. Nevertheless, the change in tree ring carbon isotopic values resulting from global industrialisation is far greater than anything observed in the tree ring isotope record during the last 1100 years. The combination of a temporally stable proxy preserved within a precisely constrained chronology for the tree rings represents an opportunity to constrain a boundary to the Anthropocene.

### 3.9.3 Sulphur concentrations and isotopic ratios

Bulk analyses of the conifer *Abies alba* from near the Ernesto cave site, NE Italy, show that S concentrations within annual rings reflect atmospheric SO₂ pollution (Figure 32b), with significantly higher values in the second half of the 20th century (Fairchild et al. 2009). Both the trees and the nearby stalagmite record the overall increasing trend in the 20th century, from the 1960s onwards, with the trees showing a more immediate (earlier) response to the
high S emissions of the 1960s and 1970s (Fairchild et al. 2009, Wynn et al. 2014). The *Abies* samples show considerable noise, including variation between individual trees (Fairchild et al. 2009). Furthermore, the S is mainly sourced from soil-waters rather than directly from the atmosphere, and the S is stored both in soil and biomass, which means that the original atmospheric signal is necessarily modified (Fairchild et al. 2009). These trees also show a trend in δ³⁴S towards lighter isotopes in recent rings, with pre-industrial values of +7.5‰ and modern values of +0.7‰, with no significant signature corresponding to the mid-20th century (Wynn et al. 2014, Figure 32b).

δ³⁴S is potentially a useful chronological proxy for past atmospheric S pollution, particularly as there is limited fractionation during incorporation into the tree (Kawamura et al. 2006). Three conifers studied from Kyushu Island, Japan provide a significant record from 1945 CE, where a dominant organically-bound sulphur (OBS) fraction and a minor water-soluble sulphur (WSS) fraction were discriminated (Kawamura et al. 2006). These showed a general increase, then decrease, in the second part of the 20th century, with higher values and a greater range seen in urban as compared with rural areas. Increased atmospheric S emissions from petroleum combustion with negative δ³⁴S values were mainly sourced from the Middle East since the 1950s, until legislation was introduced in 1968 CE. There is a ~5 year delay in δ³⁴S_OBS tree ring values compared with maximum atmospheric S concentrations, probably reflecting the time taken for sulphur deposited in soil to mobilise to the roots, and for the metabolism of sulphur after absorption (Kawamura et al. 2006).
<figure><img src="image_49.png" /><figcaption>**Figure 32.** a) Tree rings (*Pinus sylvestris*) samples from 1960 to 2003 CE at Niepołomice (Poland) showing changes of radiocarbon concentration (from Rakowski et al. 2013) compared with Northern Hemisphere (Zone 1) atmospheric values (from Hua and Barbetti 2004); b) European S emissions; c) S concentrations; and d) δ³⁴S in *Abies alba* from NE Italy compared with European S emissions (Wynn et al. 2014).</figcaption></figure>
### 3.9.4 The radiocarbon bomb spike and heavy metal concentrations

Atmospheric $\Delta^{14}$C decreased by about 20.0‰ between 1890 and 1950 CE in response to the Suess effect of adding low-activity fossil carbon, punctuated by a marked bomb spike in response to nuclear weapons testing, peaking in ~1963 CE that led to a level two times higher than that of pre-1850 CE times (Rakowski et al. 2013). Annual rings of a pine tree (Pinus sylvestris) taken in the Niepołomice area, near Kraków, Poland from 1960 to 2003 CE show a clear $\Delta^{14}$C bomb-spike in 1964 CE (Rakowski et al. 2013, Figure 32a). A suitable GSSP candidate should have a data set of both $\Delta^{14}$C and $\delta^{13}$C extending back to pre-industrial times; the data in the tree analysed by Rakowski et al. (2013) began close to the peak bomb-spike interval, so lack an adequate record of the transition across a potential Holocene–Anthropocene boundary. $^{239+240}$Pu records in tree rings are less studied, but the distribution of Pu in tree rings and in lake sediment cores seems to show similar trends (Mahara and Kudo 1995) suggesting Pu as a potentially suitable signal.

Historical changes in trace metal levels can be recognised in some tree ring records, although there may well be a time lag between metal deposition and the passage through soils and root system. Element concentrations in tree rings for a given year may differ markedly between trees from a single location (Watmough 1999). Studies on the sugar maple (Acer saccharum) suggest that there is both a rapid uptake of trace metals (e.g. Cu, Ni, Cr, Zn, Cd, Co and As) and a minimal lateral movement between rings (Watmough 1999), although radial movement of trace elements has been reported in other studies (Hietz et al. 2014).
## 4 Summary

Many of the signals described above are highly resolved, widespread and correlatable. Notably, heavy metals, δ¹³C and radiogenic fallout signals tend to be recorded in most environments that might conceivably be suitable for hosting a GSSP (Table 2). Other airborne signals, such as fly ash, nitrates and δ¹⁵N, and to a lesser extent sulphur and sulphates, CO₂ and CH₄ concentrations and δ¹⁸O occur in many environments and provide additional means of correlation. Biotic assemblages have been demonstrated to respond rapidly to local environmental changes and can provide useful secondary and local markers, but seem on current evidence unlikely to provide a globally correlatable and synchronous marker that is as temporally well-resolved as δ¹³C and radiogenic fallout or other airborne signals.

These more nearly synchronous and widespread signals, combined with the possibility of directly dating successions by such means as lamina counting, show that a well-chosen GSSP level might be precisely dated to the nearest year (or season) and might be correlated to other well-studied high-resolution sections with a resolution of a year or a few years. In less well-studied sections, or those lacking annual lamination, use of the very wide range of available proxy signals should still be able to provide correlative resolution of a few to several years, sufficient to widely locate a practically effective Holocene–Anthropocene boundary, even in apparently unpromising geological settings. For instance, Smith et al. (2016) identified what would be regarded as an Anthropocene succession of sub-ice shelf deposits at Pine Island in West Antarctica, by identifying bomb-derived plutonium within the deposits.

The location of such a Holocene–Anthropocene boundary within relatively recent successions – of the last few centuries – we regard as having use beyond the formal value of
the establishment of chronostratigraphic divisions. An extensive, variably interconnected
suite of physical, chemical and biological changes to the Earth System has been taking place,
some local, some global, many of which leave stratigraphically detectable traces. These
include phenomena such as the spread of various kinds of industrially related pollutants and
their biotic effects, changes to the atmosphere and hence to the Earth's radiative properties,
a complex pattern of species extinctions, extirpations and invasions, and the emergence and
evolution of technological systems on a planetary scale. For practical analysis of these
changes – many of which are ongoing, or accelerating – placing them in an effective
framework of space and time is a prerequisite. We suggest that a chronostratigraphic
Holocene–Anthropocene boundary could make a significant contribution to this analytical
process, not least in that this time interval is emerging as a critical transition in the geological
history of this planet (e.g. Steffen et al. 2016). The survey we provide here suggests that such
a boundary would be workable in practice, across the range of the environments we survey
here and likely more widely. Table 13 surveys the range of chemostratigraphic signals
recorded here and provides their onset and peak date for key environments. This shows the
wide range of potential proxies that could be used to mark a mid-20th century boundary.
<table><thead><tr><th>Environment</th><th>Black Carbon</th><th>Fly ash</th><th>Lead</th><th>PCBs</th><th>Pesticides (DDT)</th><th>NO<sub>3</sub><sup>-</sup></th><th>δ<sup>15</sup>N</th><th>Sulphur</th><th>SO<sub>4</sub><sup>2-</sup></th><th>CO<sub>2</sub></th><th>CH<sub>4</sub></th><th>δ<sup>13</sup>C</th><th>δ<sup>18</sup>O</th><th>D &/or dust</th><th><sup>239</sup>Pu</th><th><sup>14</sup>C</th></tr></thead><tbody><tr><td>Marine anoxic basins (1)</td><td>?</td><td>?</td><td>1965 (1970s)</td><td>1945 (1967)</td><td>1952 (1967)</td><td>?</td><td>x</td><td>x</td><td>x</td><td>x</td><td>x</td><td>?</td><td>?</td><td>x</td><td>1950-54 (1970s)</td><td>?</td></tr><tr><td>Coral bioherms & marine bivalves (2) (3)</td><td>?</td><td>?</td><td>Late 1940s (1970s) (2)</td><td>?</td><td>x</td><td>~1970 (3)</td><td>~1950 (3)</td><td>x</td><td>x</td><td>x</td><td>x</td><td>1955 (2)</td><td>~1950 (2)</td><td>x</td><td>1952 (1964) (2)</td><td>1958 (1972) (3)</td></tr><tr><td>Estuaries/ Deltas (4) (5)</td><td>?</td><td>1900s (5)</td><td>(1915) (4)</td><td>1950s (1965-1977) (4)</td><td>?</td><td>?</td><td>x</td><td>x</td><td>x</td><td>x</td><td>x</td><td>?</td><td>?</td><td>x</td><td>?</td><td>?</td></tr><tr><td>Lakes (6) (7) (8) (9) (10)</td><td>1940s-1950s (2004-6) (8)</td><td>~1950 (1970-1990)</td><td>1960s (Post-2000) (8)</td><td>~1960 (9)</td><td>1950s (9)</td><td>?</td><td>~1950 (7)</td><td>~1950 (10)</td><td>x</td><td>x</td><td>x</td><td>?</td><td>?</td><td>x</td><td>~1955 (1964) (6)</td><td>?</td></tr><tr><td>Peat & peat-lands (11) (12)</td><td>x</td><td>1950s (1970s) (11)</td><td>1810 (1979) (12)</td><td>(1960-1976) (12)</td><td>x</td><td>x</td><td>x</td><td>x</td><td>x</td><td>x</td><td>x</td><td>?</td><td>x</td><td>x</td><td>x</td><td>x</td></tr><tr><td>Ice (13) (14)</td><td>~1880 (~1910) (14)</td><td>?</td><td>~1940 (~1970) (14)</td><td>?</td><td>?</td><td>~1950 (14)</td><td>~1950 (1980) (14)</td><td>~1940 (~1970) (14)</td><td>~1940 (14)</td><td>~1950 (13)</td><td>~1950 (13)</td><td>1955 (13)</td><td>1980s (2010) (14)</td><td>~1850 (13)</td><td>1953 (1962) (13)</td><td>1954 (1966) (13)</td></tr><tr><td>Speleothems (15)</td><td>x</td><td>x</td><td>?</td><td>?</td><td>?</td><td>?</td><td>?</td><td>1980 (2000)</td><td>1880 (15)</td><td>x</td><td>x</td><td>~1840 (1860)</td><td>1875 (1940)</td><td>x</td><td>x</td><td>~1965 (~1975)</td></tr><tr><td>Trees (16) (17) (18)</td><td>x</td><td>x</td><td>?</td><td>?</td><td>?</td><td>?</td><td>1945 (18)</td><td>1960s (15)</td><td>?</td><td>x</td><td>x</td><td>1940 (1750) (16)</td><td>?</td><td>1960</td><td>?</td><td>1956 (1964) (17)</td></tr></tbody></table>
**Table 13.** Summary of key mid-20th century proxy signals and potential palaeoenvironments for a GSSP. The initial date represents the marked onset of the signal; the peak signal is shown in brackets. Key signals and environments are shown in boxes. Reference numbers relate to main geographical locations mentioned in the text: (1) Santa Barbara, USA; (2) Caribbean; (3) Nova Scotia, Canada; (4) Clyde Estuary, Scotland; (5) Urola Estuary, Spain; (6) Lake Victoria, Australia; (7) North America; (8) China; (9) Lochnagar, Scotland; (10) Lilla Öresjön, Sweden; (11) Northern England; (12) Switzerland; (13) Antarctica; (14) Greenland; (15) Ernesto Cave, Italy; (16) Fenno-Scandinavia; (17) Poland; (18) Ontario, Canada.

With regard to the environments most suitable for seeking a candidate GSSP, we offer the
following guidance:

* Anthropogenic deposits are common in terrestrial environments. They may be relatively thick (up to tens of metres), contain signals that closely reflect human influence, and are readily datable, particularly using technofossils. They tend to lack lateral continuity, may contain omission surfaces, are not annually laminated, and may lack 'preservability', all of which are likely to discount many such deposits as hosts for primary or auxiliary GSSP candidates. However, much the same criticisms also apply to many Holocene deposits. Suitable potential GSSP candidates may well exist within anthropogenic deposits, or perhaps at the interface between anthropogenic and non-anthropogenic strata, although none have yet been identified or described. More work is required in this area of investigation.

* Marine anoxic basins tend to display undisturbed annually resolved laminae that can be independently dated. The radionuclide fallout signal may show some smearing due to processes affecting the settling of radioisotopes through the water column, which can inhibit development of a sharp bomb spike. Even so, the early 1950s start of the signal (Table 13) is
typically robust. In these environments, heavy metals including Pb, organic compounds, such as PCBs and pesticides (Table 13), microplastics and biotic signals, including planktonic foraminifera and diatoms (and their stable isotope patterns), tend to be controlled by local environmental variations. Although they broadly display changes in the mid- to late-20th century, the onset of their signal can vary by decades between regions. Low sedimentation rates may reduce the thickness of Anthropocene strata to a few centimetres. Even so, the signals of change may still be clearly resolved and laterally extensive. This environment has potential for hosting an Anthropocene GSSP, provided that the fallout signal, especially $^{239}$Pu, is not significantly smeared by slow settling rates or by bioturbation.

* Coral growth bands can exhibit very-high resolution records, which, along with rapid growth rates commonly exceeding sediment accumulation rates, provide a potentially suitable medium for hosting a GSSP. The onset of the $^{239}$Pu signal is particularly robust as an early 1950s marker (Table 13), with no concerns about the influence of settling rates through the water column; peak signals are more variable between localities. However, reefs that are distal to nuclear detonation test sites might be more suitable so as to more clearly show the global pattern, rather than including some residues of the early, pre-1952 CE tests that only left a local signal. The Caribbean, Red Sea, western and northern Indian Ocean are distant from the main testing grounds (Waters et al. 2015) and are more likely to most clearly show the global atmospheric signal. Radiocarbon tends to have latitudinal diachroneity and shows a delayed signal in shallow-water corals of about a decade compared with atmospheric $^{14}$C and the $^{239}$Pu signal (Table 13), with increased lags also associated with greater water depth of coral growth. The decline of $\delta^{13}$C in corals is most marked in Atlantic corals from 1957 CE (Table 13), with the accumulation rate being more variable and of much smaller magnitude in the Indian and Pacific oceans. $\delta^{15}$N values are complex, related to source areas of ocean
currents and types of anthropogenic sources. Heavy metals can be concentrated within
certain species of coral, but they tend to record only local contamination events.
Temperature records are a mix of local and global factors and a single core would probably
not be suitable as a GSSP. If corals are to be considered as a potential host for a GSSP,
examples from the Caribbean, using <sup>239</sup>Pu as the primary signal, may be most suitable.

* Estuaries and deltas are typically responsive to anthropogenic influence, whether biotic change in response to environmental modifications and widespread introduction of neobiota, changes to sediment flux through alteration of catchment erosion rates or impounding sediments through dam construction, or through contamination by fly ash, heavy metals such as Pb, organic chemicals such as PCBs (Table 13), or microplastics. These environments, while offering good potential for hosting a candidate GSSP, may suffer from strong modulation by local influences, a lack of lateral continuity, and the common presence of omission surfaces.

* Lakes, despite being laterally disconnected, are found across large parts of the planet and, although commonly strongly modulated by local processes, also display numerous globally coherent signals. As with estuaries and deltas, lakes tend to be responsive to anthropogenic influence, resulting in a spectrum of highly resolved signals, including black carbon, fly ash, organic compounds, δ<sup>15</sup>N, S and <sup>239</sup>Pu which provide potential 1950s markers (Table 13). Despite the Anthropocene successions being typically thin, shallow perennial lakes are advantageous for hosting a potential GSSP given that settling rates of radiogenic fallout and heavy metals will be minimal and at the same time there should be few omission surfaces. The latter issue limits the suitability of many arid saline lakes, but numerous Northern Hemisphere meromictic lakes could host suitable candidate GSSP or auxiliary
sections. Lake sediments have been accepted as hosts for global auxiliary stratotypes for the base of the Holocene.

* Ombrotrophic peat bogs provide widespread local archives that faithfully record aerosol distribution from local, regional and global sources. The range of potential suitable signals are diverse (Table 8), with fly ash providing the most robust 1950s marker (Table 13). The behaviour of trace elements may be influenced by changing redox conditions and has to be carefully documented. As there is no varved sedimentation, and as some signals such as the Pu bomb spike show redistribution effects, hosting a primary GSSP in peat bogs seems not appropriate, but this environment may provide a useful auxiliary GSSP section.

* Glacial ice provides a broad spectrum of annually resolvable atmospheric signals including radionuclides, δ¹⁸O, sulphates, nitrates, δ¹⁵N, Pb and other metals, and so may provide a suitable mid-20th century marker (Table 13). Given the precedent of locating the Holocene GSSP in glacial ice, there is a strong argument for locating in an ice sheet at least an auxiliary stratotype GSSP section for the base of the Anthropocene, not least to show how its signals contrast with those at the base of the Holocene. A significant issue for the temporal resolution required for the Anthropocene is the significant time lag between the age of the ice and the included air bubbles. This lag affects CO₂, CH₄, δ¹³C and N₂O values, which are consistently younger than the enclosing ice. This effect is most pronounced where snow accumulation is slowest, so is typically more of an issue in Antarctica than Greenland. Even so, it has been recently demonstrated that the age of air bubbles can be matched to that of the associated ice through the use of nitrogen isotopes (Parrenin et al. 2013), which may circumvent what until recently appeared to be a significant problem. Slowly accumulating Antarctic ice tends to show only the more global signals, whereas Greenland
ice can reveal many regional Northern Hemisphere influences. Nevertheless, Antarctic
coastal ice cores do accumulate at high rates (e.g. Law Dome) and may provide an
appropriate GSSP candidate section. Care would have to be taken to locate such a
prospective GSSP in a region where the danger of calving of significant areas of ice from ice
shelves and ice loss due to global climate change is less imminent, given the shallow depths
at which the signal would be evident. Ellsworth Land, at the base of the Antarctic Peninsula
might provide a suitable location, since the observed increase in snow accumulation there
during the late 20th century is unprecedented in the context of the past 300 years (Thomas
et al. 2015).

* Speleothems suffer from modulation of environmental signals by the soil ecosystem, making the interpretation of signal variation complex. Sulphur loading, the depletion of δ³⁴S and the ¹⁴C (but importantly not ²³⁹Pu) bomb-spike can be clearly resolved in speleothems, with an onset close to the atmospheric signal, although the peak tends to show a decadal-scale lag (Table 13). These environments also tend to lack mid-20th century δ¹³C, δ¹⁸O and heavy metal shifts as potential environmental markers.

* Tree rings, compared with speleothems, are more responsive to and provide a clearer archive of changing atmospheric chemistries. They benefit from a global distribution and are capable of precise calendrical dating through dendrochronology without reliance upon external dating methods. Stable carbon isotopes provide the strongest “candidate” measure from tree rings as they have been demonstrated to record worldwide the change in atmospheric δ¹³C since the onset of global industrialisation. Such a uniformly expressed, well-characterised ~1940 CE departure (Table 13) has not been observed in the older tree-ring record, and the co-recorded chronology enables this trend to be clearly recognised.
Sulphur loading and depletion of δ³⁴S is evident in developed countries, starting in the 1960s, with enrichment of excess ¹⁴C from the early 1950s to a 1964 CE peak (Table 13), but this signal is a terrestrial marker only. Although certainly widespread, tree-ring signals can be influenced by complex local environmental processes, yet these may be resolved through a capacity for large-scale independent replication. There is no precedent for using a living or deceased organism (such as a tree) as the host for a GSSP, but as an archive containing a signal of atmospheric carbon isotope variability and precise timescale there are close parallels to the use of the North Greenland ice core (NGRIP) for establishing the base of the Holocene.

To summarize, the information provided here is a first overall appraisal of suitable palaeoenvironments as regards their overall potentials for hosting an Anthropocene GSSP and a range of auxiliary sections. Clearly, many candidates that are geologically suitable exist, and are found in a range of different depositional settings. The review presented here is a critical prelude to help guide the necessary next steps, which involve selecting a small number of maximally suitable sites from the most appropriate geological facies, and subjecting them to multi-proxy analysis in order to assess changes in stratigraphically-relevant proxies from pre-industrial times through to the present.

## 5 Conclusions

* Many widespread, correlatable, and highly resolved signals could be used as the primary marker and as secondary markers for the base of an Anthropocene Series. δ¹³C and radionuclide fallout signals are applicable across most environments. The latter provides the most abrupt signal, with a marked upturn in abundance of
radioisotopes of $^{239}$Pu or $^{14}$C, in 1952 and 1954 CE respectively, appearing to provide
a consistent horizon for correlation, whereas the peak ‘bomb-signals’ appear
diachronous.

* Varved deposits, such as those accumulating in marine anoxic basins, meromictic lakes, and estuaries and deltas, along with layers in glacial ice and growth rings, such as developed in corals, marine bivalves, speleothems and trees, represent those palaeoenvironments yielding highest resolution ‘stratigraphy’ at annual or even seasonal scale, and are hence preferred targets for searching for a likely GSSP section. The use of lamina or layer counting to determine ages can be verified, precisely in the case of tree rings, or radiometrically via $^{210}$Pb (marine anoxic basins, corals and bivalves, lakes, ice) and $^{14}$C (marine anoxic basins, corals and bivalves, lakes, speleothems and trees) dating.

* Ombrotrophic peat bogs and some anthropogenic deposits provide clear expression of numerous key stratigraphical markers and can show high stratal accumulation rates, but lack the annual lamination seen in other environments.

* Preservation of continuous laminated successions without missing ‘strata’ is essential. Estuaries and deltas are prone to missing laminae through erosive events, ice laminae from high-altitude glaciers through melting events, and speleothems through periods of aridity.

* Time lags between signal generation and recording within varves or laminae on decadal scales represents a significant correlatory problem in speleothems.

* A lag between the age of ice and the younger age of air bubbles within the ice limits the potential use of atmospheric CO₂ and CH₄ signals from ice cores, but close examination of sites of rapid accumulation on the Antarctic Peninsula may be
suitable, as well as sites in other coastal regions (Law Dome in Antarctica), or in
Greenland).

* Deep marine environments can display a significant settling delay that affects radioisotopes and heavy metal signals, though this effect mainly delays the peak signal rather than onset. Except for some marine anoxic basins and perhaps cold-water corals, the rates of sedimentation are too low and the rates of bioturbation are too high for the development of suitable sections for a GSSP.

* Proximity to the source of signals can result in local discrepancies in the timing of their initial onset within successions. In the search for a GSSP, there is advantage in looking at locations that are distant from such perturbations and provide a global signal, e.g. Antarctica for glacial ice, or the Caribbean for corals.

* Overall, the range of high-resolution stratigraphic proxy signals, and the varied and widespread nature of the palaeoenvironments where they can be systematically preserved in strata, suggests that there is excellent potential for locating an effective candidate Anthropocene GSSP together with auxiliary stratotypes.

## Acknowledgements

Lisa Barber is thanked for the drafting of Figure 5. This paper is a contribution of the Anthropocene Working Group (AWG), a constituent working group of the Subcommission on Quaternary Stratigraphy of the International Commission on Stratigraphy. Lucy Edwards and Jim Rose are thanked for their helpful reviews of the original manuscript.

## Declaration of Interest

The AWG receives no direct funding to carry out its research, and the authors declare no competing financial interests.
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