Carbon dioxide concentration in temperate climate caves and parent soils over an altitudinal...

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Carbon dioxide concentration in temperate climate caves and parent soils over an

altitudinal gradient and its influence on speleothem growth and fabrics.

Andrea Borsato1*, Silvia Frisia1, Renza Miorandi2

1 School of Environmental and Life Sciences, University of Newcastle, Callaghan

2308, NSW, Australia.

2 Gruppo Grotte S.A.T. "Emilio Roner", Corso Rosmini 53, 38068 Rovereto (TN),

Italy.

*Correspondence to: Andrea Borsato, E-mail: [email protected]

Key words: carbon dioxide, karst soils, speleothem fabrics, caves, temperature

gradient

This article has been accepted for publication and undergone full peer review but has not been through the copyediting, typesetting, pagination and proofreading process, which may lead to differences between this version and the Version of Record. Please cite this article as doi: 10.1002/esp.3706

mailto:[email protected]


ABSTRACT

Carbon dioxide concentrations in caves and parent soils in the Italian Alps

have been studied along a 2100 m altitudinal range ─ corresponding to a mean

annual temperature (MAT) range of 12°C ─ in order to investigate the relationship

between MAT, soil pCO2 and cave air pCO2, and to test the influence of soil pCO2 on

speleothem growth and fabric to ultimately gain insight into their palaeoclimatic

significance in temperate climate settings.

Our findings indicate that soil CO2 is linearly correlated to MAT and its mean

annual concentration is described by the equation: soil CO2 (ppmv) = 1112 + 460

MAT. Soil pCO2 can also be exponentially correlated to actual evapotranspiration.

The pCO2 in the aquifer is linearly correlated to MAT at the infiltration site and is

more influenced by summer soil pCO2. Cave air CO2 in the innermost part of the

caves exhibits a similar seasonal pattern, and commonly reaches concentrations of

about 15% to 35%, with respect to the corresponding soil values, and is

exponentially correlated to the MAT at the infiltration site.

The combination of these parameters (soil pCO2, dripwater pCO2 and cave air

pCO2) results in speleothem growth and controls their fabrics which are typical of

four MAT/elevation belts broadly corresponding to the present-day vegetation zones.

In the lower montane zone (100 to 800 m a.s.l.) speleothems mostly consist of

columnar fabric, in the upper montane zone (800 to 1600 m a.s.l.) both columnar and

dendritic fabrics are common, the Subalpine zone (1600 to 2200 m a.s.l.) is

characterised mostly by moonmilk deposits, whereas in the Alpine zone (above 2200

m a.s.l.) no speleothems are forming today. Therefore, fabric changes in fossil


speleothem can potentially be used to reconstruct MAT changes in temperate

climate karst areas.

INTRODUCTION

In karst regions, carbon dioxide concentration in soil and the epikarst greatly

influence landscape forms through enhancing dissolution and focusing surface flow

at infiltration points, such as fractures, sinkholes or cave entrances (Jennings, 1985;

Ford and Williams, 2007). When soil CO2 concentration persists at high levels over

long periods of time, which is typical of warm and wet tropical regions, spectacular

karst morphologies such as polygonal, tower and cone karst develop (Jennings,

1985; Ford and Williams, 2007; De Waele et al., 2009). By contrast, in mid and high

latitude karst settings, soil CO2 production diminished considerably during

Quaternary cold periods (Síbrava et al., 1986; Crowley and North, 1991) and the

geomorphological influence of soil CO2 may be outcompeted by glacial, periglacial

and fluvial processes in modelling the karst terrains (Ford, 1971, 1987; Lewin and

Woodward, 2009).

Altitudinal belts in temperate mountain regions, such as the Italian Alps, can

be considered analogues of latitudinal changes from the mid to high latitudes in the

Northern Hemisphere, and changes in CO2 production from the forested valley

bottom to the barren peaks above the treeline can provide insight into the effects that

Quaternary glacial and interglacial periods exerted on soil and epikarst CO2 levels.

Cave secondary mineral formations, known as speleothems, are one of the

most valuable resources to understand Earth surface conditions when glaciers and


ice sheets waxed and waned (Fairchild and Baker, 2012), and their formation

commonly depends on soil CO2 production. Carbon dioxide in the soil dissolves in

infiltration water forming weak carbonic acid which then dissolves carbonate from the

rock. From the aquifer, the water reaches the cave with an excess of CO2 with

respect to the cave environment which degasses to the atmosphere increasing the

calcite supersaturation of the water and, thus, promote speleothem formation

(Dreybrodt, 1988; Genty et al., 2001). The supersaturation state of the parent water

has been observed to exert considerable influence on speleothem growth and

fabrics (Frisia et al., 2000; Borsato et al., 2000; Frisia & Borsato, 2010; Frisia, 2015).

Thus, there is a link between speleothems in the karst subsurface, and soil CO2

production at the surface.

In temperate climate settings soil and cave pCO2 are usually positively

correlated with surface temperature (cf.: Ek and Gewelt, 1985; Bourges et al., 2001;

Spötl et al., 2005, Faimon and Ličbinská, 2010; Wong and Banner, 2010; Frisia et

al., 2011). Thus a systematic study of soil CO2 production and its controls on

speleothem fabrics across an altitudinal/temperature gradient should provide a

benchmark for understanding karst surface conditions during past climate changes

such as glacial/interglacial transitions and/or along a latitudinal gradient.

In this paper we aim at: i) verifying the correlation between

elevation/temperature and soil CO2 in a temperate climate setting over a 2,100 m

altitudinal gradient; ii) reconstructing the relationship between cave air CO2 and the

corresponding soil CO2 and iii) understanding how soil pCO2 influences speleothem

growth and fabrics to test their significance as resources to reconstruct regional and

global climate and environmental changes (cf. Frisia and Borsato, 2010; Frisia,

2015).


Speleothem growth and fabrics are also influenced by local factors, such as

bedrock composition, the amount of infiltration, the rock-water interaction and prior

calcite precipitation that can occur in the cave and in the aquifer above the cave. The

interplay of these variables is investigated in a parallel paper specifically dealing with

dripwater chemistry in the same karst region (Borsato et al., submitted).

Soil air CO2 concentration and fluxes

In karst regions, carbon dioxide concentration in soils controls bedrock

weathering and carbonate dissolution (Atkinson, 1977; Solomon and Cerling, 1987),

CO2 efflux to the atmosphere, and CO2 transfers in the subsurface that ultimately

result in cave development and/or speleothem precipitation (Dreybrodt, 1988).

Carbon dioxide production in soils is a product of autotrophs metabolic processes

(roots, and subordinately algae and chemolithotrophs), plant respiration, and to a

minor extent, of the respiration of heterotrophs (soil microorganisms and

macrofauna) (Kuzyakov, 2006, 2011). In temperate climates, CO2 concentrations in

karst soils are significantly higher than atmospheric values and usually range from

1,000 up to 15,000 ppm, with higher values in summer and lower in winter (Spötl et

al., 2005, Faimon and Ličbinská 2010; Frisia et al., 2011). Soil CO2 concentration is

a balance between CO2 production and CO2 losses, through molecular diffusion with

the atmosphere, via dissolution in soil water, and by chemical weathering of soil

minerals (Cerling, 1984).


Soil CO2 effluxes and concentrations along altitudinal gradients have been

investigated in several temperate climate settings: in most cases a quasi-linear

relationship exists between mean annual temperature and soil pCO2, concurring with

an inverse correlation between organic-C content and temperature (Amundson and

Davidson, 1990; Raich and Schlesinger, 1992; Quade et al., 1989; Dahlgren et al.,

1997; Egli et al., 2006). At the local scale, however, soil pCO2 is controlled by

additional processes, including soil moisture, soil diffusivity and storage, and wind

speed (Kuzyakov and Gavrichkova 2010; Kuzyakov 2011; Sanchez-Cañete et al.,

2011). In arid and semi-arid climates, in fact, soil moisture, rather than temperature,

becomes the major factor controlling CO2 production and efflux rates (Raich and

Potter, 1995; Almagro et al., 2009). Thus, empirical functions predicting soil CO2

based on actual evapotranspiration instead of temperature have been developed

(Brook et al., 1983).

Cave air CO2 concentration and cave ventilation

Natural fluxes of CO2 in caves are principally related to direct diffusion of gas

from the soil zone that has penetrated the epikarst, (Ek and Gewelt, 1985; Bourges

et al. 2001) and, to a minor extent, to dripwater degassing (Frisia et al., 2011;

Breecker et al., 2012). Additional sources are related to microbial oxidation of

organic matter in cave sediments and speleothems (Ortiz et al., 2013), animal

respiration, CO2 respired by plant roots that penetrate in the cave through fractures,

decomposition of organic matter in infiltration waters (Atkinson, 1977), and advection


of CO2 rich ground air derived from deeper levels (Mattey et al., 2010, Faimon et al.,

2012a)

Cave CO2 concentration is a balance between CO2 influx and efflux, which is

controlled by the exchange between cave air and the external atmosphere, and is

often referred to as cave ventilation. Cave ventilation is controlled by pressure and

temperature gradients that exist between the cave interior and the surface. It is also

controlled by the dimension and geometry of cave passages and by the presence of

multiple entrances (or fracture systems), and it is responsible for cave air pCO2

values commonly lower than soil levels. Fairchild and Baker (2012) summarized the

following mechanisms promoting cave ventilation: i) cave breathing, i.e. air exchange

forced by atmospheric pressure changes through narrow orifices; ii) chimney effect

circulation, which depends on temperature-driven surface to cave air density

differences, and is typical of cave systems with two or more entrances located at

different altitudes; iii) convective motion, which is the result of free thermal

convection or of forced ventilation, where a density current exists, such as in

descending or ascending caves; iv) water-induced flow, driven by the presence of

significant water flows; v) wind-induced flow.

In shallow or ventilated caves CO2 concentrations are significantly lower than

in the soil above the caves, and vary between 500 to 10,000 ppmv (Ek and Gewelt,

1985; Spötl et al., 2005; Baldini et al., 2008; Mattey et al., 2010; Kowalczk and

Froelich, 2010; Faimon and Ličbinská, 2010, Frisia et al., 2011). However, in deep,

confined part of caves, CO2 concentrations up to 40,000 ppmv (Bourges at al., 2001;

Wong and Banner, 2010) and occasionally up 70,000 ppmv have been measured

(Denis et al., 2005; Benavente et al., 2010).


Soil pCO2 controls the amount of dissolved bicarbonate in the soil water, in

the aquifer and, eventually, in cave dripwater driving the growth of speleothem

modulated by CO2 degassing in the cave. Thus, it becomes clear that the difference

between soil and cave air pCO2 controls the formation of speleothems and their

growth rates (Dreybrodt, 1988). In examples reported from temperate and tropical

climate setting, cave dripwater pCO2 does not vary considerably throughout the year,

and the fundamental driver for speleothem growth rates is the difference between

dripwater pCO2 and cave air pCO2 (Frisia et al., 2011).

A direct relationship between soil and cave pCO2 is not valid where carbonate

dissolution is enhanced by the oxidation of organic matter and/or sulfide mineral

oxidation occurring in the epikarst and in the aquifer (Galdenzi et al., 2008). Both

these processes promote carbonate dissolution and dripwater supersaturation, thus

resulting in speleothem growth in caves where soil at the surface is scarce or absent

(Vollweiler et al., 2006). Another process promoting speleothem growth in caves

where soil CO2 production is limited, is the presence of microbes adapted to low-

nutrient, calcium-rich environments which mediate the precipitation of calcite

moonmilk deposits (Borsato et al, 2000; Blyth and Frisia, 2008; Banks et al., 2010).

In this paper we will therefore, focus on “classic” case studies where karst dissolution

and speleothem growth are driven only by soil processes and in-cave degassing.

CAVE LOCATION, MORPHOLOGY AND CLIMATE SETTING

In order to assess the altitudinal gradient of pCO2 in caves and corresponding

soils, caves opening from the valley bottom (225 m a.s.l.) up to 2320 m a.s.l. in the


Italian Alps were considered (Fig. 1, Table 1). Caves with sub-horizontal to slightly

inclined passages were preferentially selected. Temperature and pCO2

measurements reported here were taken in the deeper part of the caves, in order to

avoid enhanced ventilation, which characterizes near-entrance passages and

multiple entrance galleries (Faimon et al., 2012a; Faimon and Lang, 2013).

Climate in the Trento province is humid temperate, with a marked orographic

effect influencing rainfall amount, which ranges from 911±154 mm year-1 at the valley

bottom (Arco station near Bus del Diaol) up to 1432±294 mm year-1 at 2360 m a.s.l.

(Rif. Graffer station, near Torrione di Vallesinella) (Table 1). The precipitation values

are the average for the 1961-1990 period (from www.meteotrentino.it) corrected for

wind-induced error on snow measurements (Yang et al., 1994). Snow cover usually

lasts from January to February at low elevation sites, and from November to April-

May at the higher altitudes

The rock overburden above the relevant sampling points is usually a few tens

of meters and can be calculated by considering the difference between the estimated

infiltration elevation and the cave entrance elevation (cfr. Table 1). Notable

exceptions to this approach are the complex, maze cave systems, where CO2

measurements were carried out in deeper parts of the caves, with rock overburden

respectively of 125 m (VP), 150 m (TV) and 210 m (CB).

The studied caves are cut in Early Jurassic, well-bedded, shallow marine

limestones (Calcari Grigi formation), with the exception of TV, which is entirely cut in

Late Triassic, well bedded, dolomites (Dolomia Principale). Both Calcari Grigi and

Dolomia Principale consist of >98.5% carbonate, with a small percentage of silica,

clay minerals, and phosphates, and negligible organic matter and sulphide content

(Borsato et al., 2007).

http://www.meteotrentino.it/


Vegetation zones and soil classification

Vegetation zones on carbonate parent materials in the region have been

classified by Sartori et al. (2005) as follows: i) the lower montane zone from the

valley bottom up to 800 m a.s.l., characterized by deciduous forest with thermophilic

(Fraxinus ornus, Ostyrya carpinifolia, Quercus pubescens) and mesophilic (Carpinus

betulus, Fraxinus excelsior) species. ii) the upper montane zone (800 to 1600 m

a.s.l.), characterised by mixed forest with beech (Fagus sylvatica), pine (Pinus

sylvestris) white fir (Abies alba), and spruce (Picea abies). iii) the subalpine zone

(2200 – 1600 m a.s.l.), characterized up to 1900 m a.s.l. by conifer forest with

spruce (Picea abies), larix (Larix decidua), pine (Pinus sylvestris), white fir (Abies

alba), and Swiss stone pine (Pinus cembra). From 1900 up to 2200 m the

vegetation association consists of dwarf mountain pine (Pinus mugo), scattered larix

(Larix decidua) and Swiss stone pine (Pinus cembra) and frequent shrubs such as

heather (Calluna vulgaris) and rhododendron (Rhododendron ferrugineum). iv) the

Alpine zone (3000 – 2200 m a.s.l.), characterized up to 2500 m a.s.l. by meadows

with sesleria (Sesleria albicans) and carex (Carex firma).

Soils on stabilized carbonate substrates in the Alpine and Subalpine zones

are classified as Alpine grassland rendzinas (Rendzic Leptosols), with thin

carbonate-free to poorly calcareous, neutral and humic soils up to 30-40 cm thick

(Sartori et al., 2005). Above an elevation of 2000 m a.s.l., soils are discontinuous

and patchy, forming pockets within bedrock. In the montane belt, forested rendzinas

(Hyperskeletic-Rendzic Leptosols) are common on debris talus and on steep or


eroded slopes, whereas on less steep and more stable slopes brown calcareous

soils (Calcari-Mollic Cambisols), up to 100 cm thick, dominate under dense forest

cover (Sartori et al., 2005; Egli et al., 2008).

METHODS

Spot cave air temperature measurements were made during each visit along

the main axis of the cave with a Vaisala GM70 meter equipped with hand-held

temperature probe. In DL, EB, GZ, ER, VP and CB multi-year temperature logging

was carried out by using Infralog DK-500 data loggers (resolution 0.03°C, precision ±

0.1 °C) and Optic StowAway loggers (resolution 0.15°C, precision ± 0.2 °C) (zero-

point calibrated by ice-water baths) (Miorandi et al., 2007; Miorandi et al., 2010).

Cave atmosphere CO2 concentration was measured by a Vaisala Meter GM70

equipped with GMP222 probe (accuracy at 25°C ±20 ppmv CO2). To avoid the

influence of a respired CO2 contribution to the cave atmosphere (Ek and Gewelt,

1985; Baldini et al., 2006), measurements were carried out by a single operator,

entering the cave alone, and the measurements were taken in a few tens of seconds

at the relevant points during each visit. This same approach was taken at Grotta di

Ernesto and resulted in a negligible influence (< 20 ppmv) of respired CO2

contribution during each measurement (Frisia et al., 2011).

Soil CO2 was sampled at ER and GZ through permanent soil gas samplers in

the B/C horizons (description in Frisia et al., 2011). In all the other caves, spot

measurements were carried out by a portable steel probe connected to a Draeger

Multiwarn gas analyser (precision ± 10%, resolution 0.01%) at various depths in the


B/C horizons as detailed in Table 2. At each sampling site, a trench was dug in the

soil down to the B/C interface in order to identify the horizons and classify the soil.

The pCO2 measurements were carried out from January to early March (winter

season) and from June to September (summer season). Each soil pCO2

measurement is the mean value of two or three measurements repeated in an area

of about 20 m2. All pCO2 measurements were corrected to standard atmospheric

pressure (1013 hPa) by using the following relationship: CO2 (ppm volume) = CO2

(measured) x 1013/pressure (measured). For all soil and cave pCO2 measurements,

the reported mean annual values were calculated as the average of summer and

winter values. Measurements in high elevation soils had to account for the rockiness

of the surface due to scanty soil cover and bedrock outcrops. For this reason,

rockiness (Rk in % of the average surface) was estimated for each soil site (Table 2)

and the measured CO2 corrected through the relationship:

Actual soil CO2 (ppmv) = Measured soil pCO2 · (1 - Rk) + 380 ·Rk.

Annual actual evapotranspiration (AET) values were calculated as the sum of

the AETs of each month. The monthly value of AET is equal to the monthly value of

the potential evapotranspiration (PET) calculated with the Thornthwaite method

(Thornthwaite and Mather, 1957), when precipitation (Pr) exceeds PET. When Pr <

PET, the monthly values of AET are assumed to be equal to Pr.


RESULTS AND DISCUSSION

Cave, soil and surface temperatures

In the inner part of the caves temperature is stable (annual oscillation

<±0.2°C) throughout the year and the temperature range for the considered set of

caves is about 12°C (Table 1), reflecting the mean annual temperature (MAT) range

at the surface from 12.5 to 0.6°C. It is therefore possible to compare the measured

cave temperatures with the long-term (1961-1990) average MAT from meteorological

stations in the region over the studied elevation range (Figure 2).

The regional MAT is described by the regression line:

MAT (°C) = 13.75 –5.65z (R2 = 0.97; p < 0.0001) (1)

where z is the elevation in km

The theoretical MAT at the cave entrances and at the infiltration elevation for

each site can then be calculated by equation (1), and these values can be compared

with the temperatures measured in the soil and in the deeper part of the caves

(Table 1). Most of the cave temperatures in the intermediate elevation range are

closely aligned with the theoretical MAT line. The complex wind-tube air circulation in

CB and TV resulted in slightly higher cave temperatures, whereas the strong positive

anomaly of EB (+ 2.1°C) is here related to the south exposure of the slope above the

cave and to the ascending trend of the gallery. On the other hand, the negative

temperature anomaly recorded for DL (-1.3°C) is related to the NW orientation of the

slope above the cave.


Soil temperatures were continuously measured at Grotta di Ernesto for ca. 5

years at a depth of 0.9 m. The average value over the 2005 to 2007 year period was

7.05 ±3.35°C, corresponding to the theoretical MAT at the site (7.08°C), and slightly

higher than the mean annual cave temperature for the same period (6.67 ±0.06°C).

The surface environment: altitudinal trend in soil CO2 concentration

Soil pCO2 measurements are reported in Table 2, along with the calculated

mean actual evapotranspiration (AET) values. Soil CO2 concentration along the

altitudinal gradient varies from 9369 ±1413 ppmv (DL near the valley bottom) to 1710

±775 ppmv (TV, the highest site) in summer and from 3130 ±502 (DL) to 380 ppmv

(TV, assumed value for the frozen soil) in winter.

Systematic monthly measurements of soil pCO2 were carried out only at

Grotta di Ernesto site, and revealed a strong correlation between soil pCO2 and

temperature (Frisia et al., 2011). The peak in soil CO2 values was recorded in

September and October (up to 12,000 ppmv), with a delay of up to two months with

respect to surface temperature. A sharp decrease in soil CO2 concentration from

~9500 ppmv to ~2000 ppmv occurred in October and November coinciding with a

drop in soil temperature of 2-3ºC (Frisia et al., 2011).

Soil pCO2 measurements above Cogola di Giazzera were carried out from

mid-January to early March (winter season) and from late July to early October

(summer season). The soil pCO2 shows a similar seasonal trend to that recorded at

Grotta di Ernesto, but with more subdued differences in values between the cold and

warm seasons (Table 2). We interpreted this different behaviour as related to the


sunnier exposure of the slope and Cogola di Giazzera (SW), which results in milder

winters, and hence more efficient soil activity, than at Grotta di Ernesto.

For all sites, there are clear relationships between seasonal soil pCO2 and

MAT/ elevation (Fig. 3A), which are described by the following linear equations:

Summer soil CO2 (ppmv) = 2067 + 690 T (R2 = 0.91; p < 0.001), (2)

Mean soil CO2 (ppmv) = 1112 + 460 T (R2 = 0.96; p < 0.0001), (3)

Winter soil CO2 (ppmv) = 156 + 231 T (R2 = 0.98; p < 0.0001), (4)

where T is the MAT in °C.

The relationship between summer soil pCO2 and elevation is best described

by a 2nd order polynomial. This non-linearity can be reasonably explained by the

observed reduced infiltration at the low altitude DL site due to a combination of high

temperature and low precipitation in summer that results in slightly lower CO2

production in the soil and promotes CO2 diffusion from soil gas to the atmosphere.

This interpretation is supported by the relationship between the summer soil pCO2

and the actual evapotranspiration (Fig. 3B), which is described by the following

exponential equation:

Summer soil CO2 (ppmv) = 380 e0.054AET (R2 = 0.92; p < 0.0001), (5)

where AET is in mm

In Fig. 3B the exponential equation proposed by Brook et al. (1983) for the

global world model is also shown for comparison:


Summer soil CO2 (log ppmv) = −3.47 + 2.09 (1−e−0·0172 AET), (6)

The discrepancy observed between equations (5) and (6) is significant only

for high AET values and suggests that summer soil CO2 can be potentially modelled

from actual evapotranspiration data.

At the same time, in temperate regions with annual precipitation exceeding

1000 mm, the relationships between MAT and CO2 concentration described by

equations 2, 3 and 4 can be successfully utilised, although it has to be noticed that

these equations are site-specific and valid for the same boundary conditions, i.e.

within the same regional climate setting, and have to be calculated for each karst

area by taking into account the influence of the local microclimate.

The cave environment: altitudinal trend in cave air pCO2

CO2 concentration in cave air varies from 10725 ±377 (DL) and 414 ±42 ppmv

(TV) in summer and decreases from 5250 ±3106 to 364 ±28 ppmv respectively in

winter (Table 3). The seasonal variability in cave air pCO2 at Grotta di Ernesto is

correlated with soil air pCO2, with a rapid increase from March to June and an abrupt

decrease from October to November (Frisia et al., 2011). As soil air pCO2 is strongly

correlated with soil temperature, and, eventually, with surface air temperature, a

clear correlation exists between cave air pCO2 and surface air temperature (Fig. 4).

The comparison between cave air pCO2 and surface temperatures at Cogola di

Giazzera provide robustness to the correlation between the two variables (Fig. 4),

although in this case cave air pCO2 has a delay of about one month with respect to

the surface temperature between March and July. As observed for ER site, this may


reflect temperature inertia in the soil related to snow cover and to the snow melt,

which usually occurs in April.

The common seasonal trend in cave air pCO2 recorded at ER and GZ sites

suggests that seasonal “spot measurements” carried out in the other caves provide a

reliable estimate of the maximum (summer season) and minimum (winter season)

cave air pCO2 values.

Summer, winter, and mean annual values (calculated as the arithmetic

averages of summer and winter values) of cave air pCO2 are correlated with surface

mean annual temperature ─ and inversely correlated with the infiltration elevation ─

through the following exponential equations (Fig. 5A and B), which take into account

the mean seasonal external atmospheric pCO2 values:

Summer cave air CO2 (ppmv) = 362 + 51.498 e0.4315Ti (R2 = 0.98; p < 0.0001), (7)

Winter cave air CO2 (ppmv) = 356 + 6.7127 e0.5052Ti (R2 = 0.95; p < 0.0001), (8)

Annual cave air CO2 (ppmv) = 359 + 28.594 e0.4471Ti (R2 = 0.97; p < 0.0001), (9)

where Ti is the MAT at the infiltration elevation in °C.

Given the small difference between infiltration elevation and cave entrance

elevation a very good correlation also exists between cave air CO2 and the MAT at

the cave entrance elevation (cf. Table 1), with R2 varying from 0.96 to 0.98. On the

other hand, the correlation with the MAT measured inside the caves is weaker (R2

varying from 0.88 to 0.91), as cave temperature is controlled by several factors,

including cave depth, morphology, fluid flow and regional geothermal gradient

(Fairchild and Baker, 2012).


The length of the caves clearly influences their air pCO2. In all the smaller

caves consisting of a single gallery (GZ, ER, EB), cave air pCO2 tends to plot along

or below the exponential fit lines, whereas in the longer caves (VP, DL) it tends to

plot above. Bus del Diaol (DL) exhibits a large excess of cave air pCO2 with respect

to the exponential fit, especially during the winter season (Fig 5A), and a higher cave

air pCO2 with respect to the corresponding soil pCO2 (cf. Table 2). Cave air pCO2

higher than soil values has been documented in other mid latitude caves (Renault,

1968; Faimon et al., 2012b): in a study of Aven d’Orgnac, Bourges et al. (2001)

demonstrated that high cave air CO2 concentrations (values up to 45,000 ppmv) are

related to to poor ventilation and CO2 fluxes from the soil and epikarst entering into

the cave through a microfissure network. This interpretation can also be applied for

Bus del Diaol as supported by several dripwater analyses which exhibit a mean

dissolved CO2 concentration corresponding to 7290 ±2475 ppmv with maximum

values of 11800 ppmv (Borsato et al., 2007; Borsato et al., submitted). The dripwater

values are systematically higher than the corresponding soil and cave air values,

suggesting additional CO2 generation in the epikarst possibly related to the presence

of deep soil pockets where organic matter has accumulated (Atkinson, 1977; Faimon

et al., 2012b). Deep soil pockets or aquifer storage, would result in release of CO2 to

the cave air even during cold and dry seasons, when the soil zone at shallower

depth is affected by reduced infiltration and CO2 production.

Despite the singularity of Bus del Diaol, the exponential correlations between

cave air CO2, and MAT (eq. 7 to 9) are robust, and describe regional environmental

conditions that can be extended to other karst areas in temperate climate settings.


Relationships between soil, aquifer and cave air pCO2

The results of the soil (Fig. 3A) and cave air (Fig. 5A) CO2 concentrations at

different infiltration elevations are plotted in a single graph in Figure 6 to highlight

their connection and seasonal trends. The curves show an overall antipathetic trend,

which clearly indicates that there is no direct diffusion from the soil to the caves

through the epikarst. Critically, studies conducted at Grotta di Ernesto demonstrated

that strong seasonal fluctuations in soil pCO2, are not reflected by similar variations

in the aquifer, where pCO2 remains relatively constant throughout the year. This

phenomenon is related to the filling of the aquifer by water, whose residence time

has been estimated to be of the order of a few months. Filling of the aquifer on the

one hand may cut off the flux of CO2 from the soil, on the other hand results in

mixing of soil and stored CO2 (Miorandi et al., 2010; Frisia et al., 2011). Such role of

the epikarst and aquifer is not exclusive to ER but it is common for most of the

studied caves (Borsato et al., 2007; Borsato et al., submitted) and in other case

studies (Fernandez-Cortez et al., 2009).

By considering that most of the infiltration occurs between spring and autumn

(Miorandi et al., 2010), the actual aquifer CO2 concentration probably reflects the

spring-to-autumn soil pCO2, i.e. values intermediate between the summer and the

annual mean values. This assumption is corroborated by measurements conducted

at Grotta di Ernesto, where the less evolved dripwaters (less modified by in-cave

degassing) have a composition intermediate between the summer and the annual

mean soil pCO2 values (Fairchild et al., 2000; Frisia et al., 2011). Therefore, the

karst aquifer acts as a CO2 reservoir, which incorporates preferentially summer soil

pCO2, whereas winter soil gas does not contribute significantly to the aquifer pCO2.


In the case of Bus del Diaol, where cave CO2 exceeds soil CO2 concentration,

the measured soil pCO2 values are not expected to be representative of the pCO2 in

the aquifer, and a mean pCO2 value equal, or higher than the corresponding cave air

pCO2 should be expected for aquifer waters. If we assume that the actual DL “soil”

values (which include the contribution from CO2 evolved in deep soil pockets) equal

the measured cave air values (cf. Table 3), the soil CO2, as well as the aquifer CO2

values, can now be fitted by linear equations (Fig. 7):

Summer soil CO2 (ppmv) = 1650 + 814 Ti (R2 = 0.97; p < 0.0001), (10)

Mean annual soil CO2 (ppmv) = 667 + 432 Ti (R2 = 0.97; p < 0.0001), (11)

Mean annual aquifer CO2 (ppmv) = 1041 + 741 Ti (R2 = 0.98; p < 0.0001), (12)

where Ti is the MAT at the infiltration elevation in °C.

Figure 7 shows the relationships between cave air CO2, soil CO2 and aquifer

CO2 concentrations with respect to MAT and the infiltration elevation. The difference

between aquifer and cave air pCO2, highlights the potential for dripwater degassing

to the cave atmosphere: a high value of potential degassing characterises the winter

season, whereas in summer degassing is less efficient, especially in caves in the

low- and high- elevation ranges. This has obvious implications for the

supersaturation state with respect to calcite of cave waters and, consequently,

underground karst processes.


IMPLICATIONS FOR SPELEOTHEMS AS PALAEOCLIMATE ARCHIVES

Speleothem growth has been found to be dependent on the dripwater Ca2+

content and CO2 degassing in the cave which control the supersaturation state of the

parent water, (Baker et al., 1998; Kaufmann, 2003; Kaufmann and Dreybrodt, 2004)

and, therefore, is possibly related to surface MAT (Dreybrodt, 1988; Baker et al.,

1998) (cf. Fig. 3). There has been much effort to extract past temperature variations

from the oxygen isotope ratios of speleothem carbonate and fluid inclusions, whilst

temperature-related soil and vegetation processes have been reconstructed from the

C isotope ratios in the carbonate (cf. Fairchild and Baker, 2012). The physical

characteristics of cave stalagmites and flowstones have been recognized in

speleothem palaeoclimate research as unique archives of the evolution of the

environment through their direct link with vegetation cover and karst hydrology

(Frisia, 2015). The combination of temperature effect on soil CO2 production, the

transmission of the soil CO2 to dripwater via the aquifer, and, finally, the cave air

pCO2 control the development of stalagmite fabrics by modulating the

supersaturation state of the parent water with respect to calcite (Frisia, 2015).

In figure 7 the relationships between cave air CO2, soil CO2 and aquifer CO2

concentrations with respect to MAT and infiltration elevation is compared to the

dripwater calcite saturation index and to the distribution of speleothem fabrics

(Borsato et al., 2000; Frisia et al., 2000; Borsato et al., 2007; Frisia and Borsato,

2010; Frisia, 2015) and is complemented by the relationship between MAT and

modern speleothem δ13C, which has been documented for caves in the same karst

region (Johnston et al., 2013). This trend is directly related to the higher proportion of

biogenic (light) δ13C at lower elevation, whereas at high altitudes the δ13C signal is


closer to atmospheric values (Fohlmeister et al., 2011; Frisia et al., 2011). Therefore,

speleothem fabrics coupled with δ13C analyses could be utilised in paleoclimate

reconstructions to obtain information on past soil pCO2 and soil MAT.

The distribution of speleothem fabrics in the studied caves allows the

identification of four elevation zones broadly corresponding to the present-day

vegetation zones and characterized by their temperature-dependant soil, aquifer and

cave pCO2. These fabrics, and their C isotope signal, can then be potentially used

for the “backward modelling” of past temperature-related karst surface processes.

● Lower montane zone (elevation: 100 to 800 m a.s.l.; MAT range: 13.2 to 9.2°C).

The zone is characterized by high pCO2 in soil and epikarst, which promotes

rock carbonate dissolution, high dissolved Ca2+ content and dripwater with the

highest supersaturation of the whole altitudinal range, with mean calcite SI from 0.5

up to 1.0 (Frisia, 2015). Stalagmites, the most common speleothem used for

palaeoclimate reconstructions, consist of calcite crystals, arranged in compact

columnar fabric in poor ventilated parts of the caves (Frisia and Borsato, 2010;

Frisia, 2015). Contrary to theoretical predictions (Baker et al., 1998), the growth of

compact columnar fabric can be very slow, in the order of a few µm/year (Belli et al.,

2013) in cases of high cave air pCO2 such as Bus del Diaol. In more ventilated cave

passages, stalagmites may exhibit dendritic fabric, whose model of development

requires strong degassing. However, the dendritic fabric is uncommon in the lower

montane zone, where it is most typical near cave entrances, where bio-influenced

precipitation results in the development of feather dendrite crystals (cf. Jones and

Renaut, 2010). The 13C of columnar fabric commonly shows the most depleted


values, consistent with a high proportion of soil-derived CO2 in the parent water

(Johnston et al., 2013; Frisia, 2015).

● Upper montane zone (elevation: 800 to 1600 m a.s.l.; MAT range: 9.2 to 4.7°C).

In this zone, the concentration of dissolved CO2 in the aquifer typically varies

between 2500 and 5000 ppmv, which, in pure limestone aquifers, results in

moderate levels of dissolved Ca2+ (ca. 1-2 mmol L-1) and slightly supersaturated

dripwater (mean SI calcite 0.0 to 0.5). Speleothem growth is restricted to or

enhanced in the cold season, when ventilation controlled by thermal convection

removes CO2 from the cave air, and allows for a more efficient degassing of the

dripwater during the time of equilibration (Hansen et al., 2013). The fabrics change

as a result of different drip rates and supersaturation levels (Frisia et al., 2000; Frisia,

2015). Compact columnar fabrics are to be expected when relatively constant drip

rates are associated with low supersaturation values, such as the case of Grotta di

Ernesto today. Dendritic fabrics are common when variable drip rates and strong

seasonal cave air pCO2 fluctuations prevail, such as the case of many stalagmites at

Cogola di Giazzera.

It is worth mentioning that during the cooler Little Ice Age (LIA), the same

stalagmites from Grotta di Ernesto that are characterized by compact columnar

fabric in the last 150 years, exhibited columnar microcrystalline fabric dominated by

organic-rich laminae and low growth rates (Frisia et al., 2003). Because of the colder

temperature during the LIA it is reasonable to infer a lower soil and aquifer pCO2

resulting in a lower calcite supersaturation and a lower growth rate also enhanced by

the action of growth inhibitors (cf. Frisia et al., 2000).


● Subalpine zone (elevation: 1600 to 2200 m a.s.l.; MAT range: 4.7 to 1.3°C).

In this zone, low soil pCO2 results in low dissolved Ca2+ concentration in the

dripwater, which is either undersaturated or barely at saturation (mean SI calcite 0.0

to -0.3). Speleothem growth becomes rare, and related to peculiar situations such as

near entrances and highly ventilated galleries, where both evaporation and

degassing increase the mineralisation and saturation of the parent water. In this

altitudinal belt, speleothem growth is promoted by other processes, rather than

degassing alone, the most common of which is bio-influenced (or bio-induced)

carbonate precipitation in the form of moonmilk (Borsato et al., 2000). Microbial

respiration increases water supersaturation and favour the formation of elongated

calcite fibres in moonmilk deposits as documented in CB and CZ caves (Borsato et

al., 2000; Blyth and Frisia, 2008). These deposits are characterized by more positive

13C values than the lower altitude speleothems, which indicates a low soil efficiency

and a higher contribution of atmospheric CO2 to the 13C signal.

● Alpine zone (elevation above 2200 m a.s.l.; MAT below 1.3°C).

Above 2200 m a.s.l., patchy soil and the snow cover lasting over six months

lead to very low pCO2 and Ca2+ concentration in the aquifer and results in

undersaturated dripwater. Speleothems do not precipitate, unless processes such as

sulphide oxidation, as documented for the high Alpine (2500 m a.s.l.) Spannagel

cave in Tyrol (Vollweiler et al., 2006), occur. Hence, the occurrence in this altitudinal

belt of fossil speleothems consisting of columnar calcite indicates that there were

periods in the climate history of the region when the MAT at the surface was similar

to that of the upper montane, or even lower montane zones today (i.e. up to 6-7C

higher).


Fabrics as archives of temperature-related soil CO2 production

If soil CO2 production decreases/increases in response to cooling/warming,

than both the dissolved Ca2+ content and the degassing-related supersaturation

state of the drip water decrease/increase. This has been demonstrated to yield

changes in stalagmite growth rates and external morphology (Baker et al., 1998;

Genty et al., 2001; Kaufmann, 2003; Kaufmann and Dreybrodt, 2004), as well as

changes in speleothem fabrics (Frisia and Borsato, 2010; Boch et al., 2011; Belli et

al., 2013). In our conceptual model, we should expect to observe a passage from

dendritic or microcrystalline to columnar fabric for a 1-2°C increase in MAT in the

lower altitudinal range of the upper montane zone. This is the case of Grotta di

Ernesto stalagmites at the passage from the Little Ice Age to the industrial warming

(Frisia et al., 2003). By contrast, a decrease of 1-2°C in MAT occurring in the higher

altitudinal range of the upper montane zone would result in the passage from

microcrystalline or dendritic to moonmilk (cf. Frisia, 2015). Therefore, caves located

near the boundaries between different speleothem/vegetation zones are “fabric-

sensitive”, and can provide an independent proxy of temperature-related soil CO2

production.

CONCLUSIONS

In the temperate climate setting of the Southern Calcareous Alps of Italy an

almost linear relationship between elevation/MAT and CO2 concentration in soils and


aquifers, and an exponential relationship between elevation/MAT and cave air pCO2

has been illustrated here.

A conceptual framework that relates surface temperature to cave calcite

fabrics has been provided, and the elevation - temperature relationship can be

utilised as analogue for longitudinal - temperature changes in mid-high latitudes in

the Northern Hemisphere.

It is here suggested that changes in speleothem fabrics with altitude in

temperate climate settings can be used as an archive of changes in surface karst

environment, and, in particular, of temperature and related soil and aquifer pCO2

changes. The influence of surface processes in other karst regions, and in particular

in more arid and in tropical settings, should be assessed to build towards an

understanding of fabrics as archives of karst surface processes over a wider

latitudinal gradient.

ACKNOWLEDGEMENTS

The data presented in this work were acquired in the frame of the projects

AQUAPAST funded by the Autonomuos Province of Trento (Italy) and DFG

Research Group 668 DAPHNE funded by the Deutsche Forschungsgemeinschaft

(Germany). We would like to thanks Michele Zandonati for fieldwork assistance, Ian

I. Fairchild for the revision of a first draft of the paper, Jiří Faimon and two other

anonymous reviewers for the helpful suggestions and recommendations. The

Authors declare that they have no conflict of interest.


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Table 1. Location, climate and relevant characteristics of the studied caves.

1 Estimated infiltration elevation.

2 Main gallery trend: hor: subhorizontal; asc: ascending; des: descending.

3 Air mean annual temperature in the inner part of the cave.

4 Calculated mean annual temperature at the cave entrance elevation.

5 Calculated mean annual temperature at the infiltration elevation.

6 Fabrics: Col = columnar, Den = dendritic, Moon = moonmilk.

Table 2. Soil CO2 concentrations measured at and near the cave sites. For Grotta di

Ernesto the CO2 data are from Frisia et al., 2011.

1 Theoretical mean annual temperature at the site.

2 Mean depth for the soil CO2 measurements.

3 Mean annual actual evapotranspiration.

4 Frozen soil: inferred value.


Table 3. Elevation, cave morphology and ventilation and cave air CO2 concentration

for the studied caves.

1 Estimated infiltration elevation.

2 Actual elevation of the CO2 measurements.

3 Distance from the entrance of the CO2 measurements.


Figure 1. Digital elevation model of Trento province (Southern Calcareous Alps,

Northern Italy), with the location of the studied caves. Cave codes are those reported

in Table 1.


Figure 2. Air temperature of meteorological stations and cave interiors plotted

against elevation. Black dots indicate average air temperature at 2 m height over the

period 1961-1990 at 19 weather stations in the Trento province

(www.meteotrentino.it). Open diamonds represent the cave entrance elevation

plotted against the air MAT from the stable interior of the caves. Cave abbreviations

as in Table 1.


Figure 3. A) Seasonal soil CO2 measured above and near some cave sites (cf. Table

2) compared with their actual elevation and the theoretical mean annual

temperature. The linear relationships (eq. 2 to 4) with their correlation coefficients

are shown; for the summer values both the linear and the 2nd order polynomial

(dashed line) relationships are shown. The error bars for the mean year values are

not reported for clarity. B). Summer soil CO2 measurements compared with the

actual evapotranspiration. The exponential regression line (eq. 5) is compared with

the curve obtained by using Brook et al. (1983) equation.


Figure 4. Comparison between cave air CO2 concentration and surface

temperatures at Grotta di Ernesto (mean values 2005 – 2007) and Cogola di

Giazzera (mean values 2006 – 2012). The grey horizontal bars visualise the times of

the year when the external temperature exceed the cave temperature.


Figure 5. A) Mean seasonal values of cave air CO2 for the studied caves plotted

against the infiltration elevation at the surface and the corresponding MAT (cf. Table

1); the exponential fits represent equations (7) to (9). In B) the seasonal external

atmospheric CO2 values have been subtracted from the cave air CO2 values in order

to emphasise the exponential correlations.


Figure 6. Measured seasonal cave and soil air CO2 plotted against the infiltration

elevation and infiltration temperature. The regression lines correspond to eq. (2) and

(4) (soil) and to eq. (7 - 8) (cave).


Figure 7. Cave air, aquifer and soil CO2 concentrations plotted against infiltration

elevation and infiltration MAT and compared with the modern speleothem fabrics,

δ13C of modern calcite and dripwater calcite saturation index (SIcc) for the studied

caves. The soil CO2 values (soil*) for DL are corrected for the contribution of deep

soil pockets. The green vertical dashed lines define the vegetation and

elevation/temperatures zones as discussed in the text. Data for the fabric distribution

of modern forming speleothems and dripwater SIcc are from Borsato et al. (2000),

Frisia et al. (2000), Borsato et al. (2007), Frisia and Borsato (2010), and Frisia

(2015). The relationship between δ13C of modern calcite and temperature/elevation

is from Johnston et al. (2013).

Carbon dioxide concentration in temperate climate caves and parent soils over an altitudinal...

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