I think I saw that episode also. It was named “Night of the teleconnections” wasn’t it?

I’ve not read the Holdsworth paper yet. However, from recollection the Mt Logan record has both d18O and d2H at high resolution. The advanatge of this is that it is that measuring both oxygen and hydrogen isotopes allows one to calculate the deuterium excess value, d, as:

d = delta(2H) – (8 * delta(18O))

The deuterium excess value is sensitive to processes that occur in the source region and should therefore be strongly indicative of changes in either the location of the source region, or changes in processes in the source region such as relative humidity. Again from memory the deuterium excess at Mt Logan undergoes a rapid shift during the 19th century that would indicate significant changes in the source region for the water vapour.

It is to be regretted that we don’t have good deuterium data for many more ice cores and notably those at low latitude and high altitude.

Out of interest I’m just completing the oxygen isotope measurements on an ice core from the Antarctic Peninsula. These have a very high temporal resolution, better than monthly, for the past 150 years (2500+ samples). I’ve completed some preliminary deuterium measurements and have many more to do. Early indications are that the deuterium excess has remained constant.

As I mentioned, it was an old re-run. I believe that the CSI on the show re-trained and now publishes multiproxy temperature reconstructions.

There are a number of ways we can estimate the age of recharge of groundwater. Some of these are radiometric dating techniques. For young waters, say <100 years, we can use tritium. There is both a large bomb induced spike that is apparent in some groundwaters. This is as a result of atmospheric thermonuclear bomb testing in the late 1950′s and 1960′s. Whilst there has been a degree of dispersion in the spike, many aquifer systems can be modelled by piston flow and the spike still identified. This gives us a measure of flow velocity and recharge rates to an aquifer.

We can extend the tritium method using T-3He techniques. T decays to 3He and measuring both in groundwater allows explicit solution of the radiometric decay equation.

For older waters we can use 14C in the dissolved inorganic carbon load. Large corrections have to be made for dead carbon from the source aquifer but none the less reasonable dates can be determined.

Okay now for your first question. Well it is a common misunderstanding that water:rock ratios are low. Yes the pore volume to rock ratio is low, but the factor we need to consider is the volume flux of water through any unit volume of the aquifer. This can be very large in some systems. I returned from an active karst region in Ireland in September with a group of undergraduates. Here the recharge rate is about 1 m per annum. Most of this flow is focussed in a joint and fracture system with low storativity in the aquifer. This means that 1 cubic metre of water passes through every cubic metre of rock a year. In 100 years the water rock ratio is 100:1, over 1000 years it is 1000:1. It has a pH of about 4 on entering the soil zone, rapidly dissolves CO2 and then carbonate reaching saturation at elevated CO2 pressures very quickly. From then on the only possible mechanism for further exchange with the aquifer is dissolution and reprecipitation, for which there has to be a driving force (pressure solution, dissolution of high Mg calcite and aragonite and precipitation of low Mg calcite etc). or solid state diffusion in the carbonate lattice. At low temperatures solid state diffusion is out which leaves dissolution and reprecipitation.

With the large water:rock ratios that we see in these aquifers if dissolution and reprecipitation were ongoing we would see it in the oxygen isotope signature of the host carbonate which would tend towards a meteoric/freshwater value. We don’t see this. Many aquifers, despite 10′s-100′s of thousands of years and in some cases millions of years of being active still retain their original marine isotope signature in the carbonate.

He did ice cores on Mt Logan (recently updated by Fisher and others) which have more negative dO!8 in the 20th century than 19th century – contrary to the Lonnie Thompson model. They attribute this to changes in source region – an increasingly prevalent theme it seems for everyone except Lonnie Thompson.

I think that it would be prudent to keep an asterisk on the dating of Dasuopu. Some high-accumulation glaciers turn over their inventory very rapidly and have negligible information past a few centuries. Dasuopu is a high-accumulation glacier, which appears to me to have a surprisingly long time series relative to comparable glaciers (but this is a long exegesis and I’ve just dipped my toe in the data.) This would be an issue worth visiting in Thompson’s glaciers. Unfortunately Thompson has refused to provide a proper data archive so it’s a bit of a dead end.

Hu, an interesting observation wrt to Lonnie Thompson’s ice cores. The temperature effect on the isotope composition is extreme. It’s not possible to produce such effects from a simple rayleigh distillation model of an adiabatically cooling atmosphere. I would think that there is a possibility that there is a weakening amount (monsoon effect) or a change in source region in some of these cores. It may be that the amount effect is positively correlated with some climate or temperature change. However it is not the temperature change that is directly controlling the isotope composition of the precipitation.

Very interesting! Some of his cores have well-defined annual layers that might permit the rate of accumulation to be measured and controlled for. But we still wouldn’t know what the seasonality of the precipitation was.

Now for your question about the 18O composition of water and speleothems. The answer is the system is buffered by the water isotope composition. The amount of bicarbonate in solution is small compared to the amount of water. In most of these systems the water rock ratio is also high. I’m not aware of any carbonate aquifer where the rock buffers, or exerts any influence on the isotopic composition of the dissolved bicarbonate.

Its true that the amount of bicarb in solution at any moment is trivial in comparison to the amount of water, but the amount of water in limestone bedrock is also trivial in comparison to the amount of limestone. The water must have lingered in the limestone long enough to become saturated, otherwise it would be dissolving the speleothems instead of building them. But even when it’s saturated, it continues to exchange carbs and therefore Os with the limestone at the same rate as before. Isn’t it just a matter of local hydrology how long this continues and how complete it becomes?

There are systems in which the oxygen isotope composition of the groudwater/porewater is heavily influenced by the rock isotope composition. However, these are in geothermal areas.

High temperatures would accelerate the exchange, but how do we know this groundwater hasn’t been sitting around for years or even centuries?

Now for your question about the 18O composition of water and speleothems. The answer is the system is buffered by the water isotope composition. The amount of bicarbonate in solution is small compared to the amount of water. In most of these systems the water rock ratio is also high. I’m not aware of any carbonate aquifer where the rock buffers, or exerts any influence on the isotopic composition of the dissolved bicarbonate.

There are systems in which the oxygen isotope composition of the groudwater/porewater is heavily influenced by the rock isotope composition. However, these are in geothermal areas. Here the high temperatures promote water-rock interactions and the water composition changes to be in equilibrium with the aquifer at higher temeprature. There are many examples of geothermal areas where this is seen. The most significant one is the system of mid-ocean ridges where hydrothermal reaction between sea water and hot rocks exerts a significant control ohn the oceans isotopic composition.

Of course the carbon system is another story entirely. Here the only two sources of carbon are either atmospheric, organic, or the rock derived carbon.

In high latitude sites the temperature effect on rainfall and groundwater isotope composition is generally on the order of +0.4 to +0.7 parts per thousand per degree C. This is why we expect to see a net positive relationship in speleothems associated with temperate latitude caves. However, net negative relationships have been observed c.f. Mo-i-rana in Norway and some Alpine examples. This is unexpected and would seem to point towards some unexplained hydrology.

Lonnie Thompson’s 6 ice cores from his CC03 paper show a much stronger effect of temperature on d18O than this, in the 4 cases when the correlation is significant (ppk = parts per 1000, aka o/oo):

Himalayan cores:
Guliya (35.30 dN) 2.73 ppk/dC, p = .0348
Dasuopu (28.38 dN) 3.32 ppk/dC, p = .0056

Andean cores:
Quelccaya Summit (13.93 dS) 2.33 ppk/dC, p = .0364
Huascaran (9.12 dS) 1.91 ppk/dC, p = .0794 (only weakly significant)

These regressions are against HadCRU3v GL, constrained by Thompson’s archived data to be decadal. All DWs are greater than 2, so positive serial correlation is not an issue at this decadal frequency. His other two cores, Dunde (38 dN) and Sajama (18.10 dS), both have positive but insignificant coefficients less than unity.

Since the O18 in H2O and dissolved HCO3- mingles (per #77, 78), wouldn’t groundwater that has been sitting in pores in limestone bedrock for days, months, or even years have O18 that reflects that in the limestone rather than that in the source rainwater or snowmelt? Or at least be an attenuated version of the original? Could this the source of the discrepancy?